Download tethyan evolution of turkey: a plate tectonic approach

75 (1981) 181-241
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Tectonophysics,
TETHYAN EVOLUTION
APPROACH
OF TURKEY:
181
A PLATE TECTONIC
A.M. CELAL SENGGR and YCCEL YILMAZ
Department
of Geological Sciences,
State University of New
N. Y. 12222
(U.S.A.)
Jeoloji Bijliimii, Yerbilimleri
Fakiiltesi, Istanbul hiversitesi,
(Turkey)
(Received May 12,198O;
York
at Albany,
Vezneciler,
Albany,
Istanbul
revised version accepted October 20,198O)
This paper is dedicated to Professor Jan
Houghton Brunn on the occasion of his
retirement from active teaching and in
recognition of his fundamental contributions
to our understanding of the tectonics of Turkey over the past decade and-a-half.
ABSTRACT
Sengor, A.M.C. and Yilmaz, Y., 1981. Tethyan evolution
approach. Tectonophysics, 75: 181-241.
of Turkey: a plate tectonic
The Tethyan evolution of Turkey may be divided into two main phases, namely a
Palaeo-Tethyan and a Neo-Tethyan, although they partly overlap in time. The PalaeoTethyan evolution was governed by the main south-dipping (present geographic orientation) subduction zone of Palaeo-Tethys beneath northern Turkey during the PermoLiassic interval. During the Permian the entire present area of Turkey constituted a part
of the northern margin of Gondwana-Land. A marginal basin opened above the subduction zone and disrupted this margin during the early Triassic. In this paper it is called the
Karakaya marginal sea, which was already closed by earliest Jurassic times because early
Jurassic sediments unconformably overlie its deformed lithologies. The present eastern
Mediterranean and its easterly continuation into the Bitlis and Zagros oceans began opening mainly during the Carnian-Norian interval. This opening marked the birth of NeoTethys behind the Cimmerian continent which, at that time, started to separate from
northern Gondwana-Land. During the early Jurassic the Cimmerian continent internally
disintegrated behind the Palaeo-Tethyan arc constituting its northern margin and gave
birth to the northern branch of Neo-Tethys. The northern branch of Neo-Tethys included
the Intra-Pontide, Izmh-Ankara,
and the Inner Tauride oceans. With the closure of
Palaeo-Tethys during the medial Jurassic only two oceanic areas were left in Turkey: the
multi-armed northern and the relatively simpler southern branches of Neo-Tethys. The
northern branch separated the Anatolide-Tauride
platform with its long appendage, the
Bitlis-Potiirge fragment from Eurasia, whereas the southern one separated them from the
main body of Gondwana-Land. The Intra-Pontide and the IzmirAnkara
oceans isolated
a small Sakarya continent within the northern branch, which may represent an easterly
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02.50 @ 1981 Elsevier Scientific Publishing Company
182
continuation
of the Paikon Ridge of the Vardar Zone m Slacedonia.
t’he .\natolide
Tauride
platform
itself const.itutcd
the easterly
continuation
of thr ;\puliau platform
that had remained
attached
to Africa through Sicily. The Neo-Tethyan
oceans reached
their maximum
size during the early Cretaceous
in Turkey and their contraction
began
during the early late Cretaceous.
Both oceans were eliminated
mainly by north-dipping
subduction,
beneath
the Eurasian,
Sakaryan,
and the Anatolide--Tauride
margins. Suhduction beneath the Eurasian margin formed a marginal basin, the present Black Sea and
its westerly
prolongation
info the Srednogorie
province
of t,he Halkanides,
during the
medial to late Cretaceous.
This resulted
in the isolation
of a Rhodope-Pontide
fragment
(essentially
an island arc) south of the southern
margin of Eurasia. Late f:retaceous
is also
a time of widespread
ophiolite
obduetion
in Turkey,
when the Bozkir ophiolite
nappc
was abducted
onto the northern
margin of the AnatolideTauride platform.
Two other
ophiolite
nappes were emplaced
onto the Bitlis-Potiirge
fragment
and onto the northern margin of the Arabian platform
respectively.
This last event, occurred
as a result oi
the collision
of the Bitlis-Piitiirge
fragment
with Arabia. Shortiy after this tollision
dub
ing the Campanian-Maastrichtian,
a subduction
zone began consuming
the floor of the
inner Tauride ocean just to the north of the Bitlis-Poturge
fragment
producing
the arc
lithologies
of the Yiiksekova
complex.
During the Maastrichtian--Middle
Eocene interval
a marginal
basin complex,
the Maden and the qiingtig basins began opening above this
subduction
zone, disrupting
the ophiolite-laden
B‘itlis-Piitiirge
fragment.
The Anatolide--Tauride platform
collided with the Pontide arc system (Rhodope-Pontide
fragment plus
the Sakarya
continent
that collided
with the former during the latest (‘retaceous
along
the Intra Pontide suture) during the early to late Eocene interval.
This collision resulted
in the large-scale south-vergent
internal imbrication
of the platform
that produced the fartravelled nappe systems of the Taurides,
and buried beneath these, the metamorphic
axis
of Anatolia,
the Anatolides.
The Maden basin closed during the early late Eocene
by
nortll-dipping
subduction,
synthetic
to the Inner-Tauride
subc~u~~i~)~ zone that had
switched
from south-dipping
subduction
beneath
the Bitlis-F~t~rg~
fragment
to north’
dipping subduction
beneath the Anatolide-Tauride
platform
during the later Palaeocene.
Finally,
the terminal
collision
of Arabia with Eurasia in eastern Turkey eliminated
the
$Xngii~ basin as well and created
the present
tectonic
regime of Turkey
by pushing a
considerable
piece of it eastwards along the two newly-generated
transform
faulls, namely
those of North and East Anatolia.
Much of the present eastern Anatolia is underlain
by an
extensive
melange prism that accumulated
during the late Cretaceous---late
Eocene interval
north and east of the Bitlis-Piitiirge
fragment.
INTRODUCTION
The analysis of the structure of Turkey has proved ex~eptiona1~~ difficult
in the past, because of a large concentration
of a number of convergent
events throughout its history. The mountain ranges of Turkey constitute the
easternmost
segment of the Mediterranean
Alpides, where Palaeo-Tethyan
and Neo-Tethyan
deformations
have been superimposed
on a continental
basement
(partly
Pan African-Baykalian
and partly
Hercynian
ages),
whereas a portion of the Neo-Tethyan
evolution alone has been recorded by
the predominantly
late Mesozoic--Cenozoic
accretionary
complex of eastern
Anatolia (see Fig. 1). The Turkish orogen forms a critical link between the
systems, because Tethyan palaeoMediterranean
and the Asiatic Tethyan*
* The
term
“Alpide”
is used
here
to designate
the erogenic
belts
that
descended
from
183
geographic elements peculiar to each meet and terminate in the Anatolian
Peninsula (Asia Minor). Turkey is an ideal location in which to study the
relationships of the Palaeo- and the Neo-Tethyan systems, not only because
both are well-developed, but also because despite the very intense collisional
events that accompanied the closure of Neo-Tethys, pre-Neo-Tethyan-opening palaeogeography of the area can be reconstructed with some confidence
as a result of the Neo-Tethyan oceans having opened and closed roughly along
the same zones in Turkey. As Tethyan structures cross-cut Hercynian trends
in Turkey, the latter can be used as an independent check on the reconstructions. These unusually favourable conditions provide a good data base on
which to tackle a variety of strati~phic~t~ctural
problems that have so
far eluded explanation.
Because the birth and the further evolution of the Neo-Tethyan oceans in
Turkey and their tectonic relationships with older and to some extent contemporaneous st~ctures such as P~aeo-Tethys and the Cimmerian Continent (Sengiir, 1979a) can be so clearly observed and relatively easily interpreted in Turkey, the history of this segment of the Mediterranean Alpides
adds a hitherto generally unconsidered perspective to the study of the
Tethyan evolution of the Medite~ane~ region as a whole, namely an eastern,
“Tethyan” one. So far, most of the plate tectonic syntheses of peri-Mediterranean post-Palaeozoic geology have been constructed in terms of the opening history of the Atlantic Ocean (e.g. Smith, 1971; Dewey et al., 1973; BijuDuval et al., 1977, 19’78; Tapponnier, 1977; Channel et al., 1979; Burchfiel,
1980). Although this has provided a well-constrained kinematic basis for the
analysis of the tectonics of the Mediterranean area since the Toarcian, it has
done little to clarify earlier Mesozoic events that increase both in complexity
and in variety from west to east along the Tethyan chains of the peri-Mediterranean area. The “Tethyan” influence on the Mesozoic evolution of the
Mediterranean region has been at least as important as the Atlantic interference and the geology of Turkey provides an excellent data base to discover
its nature and to decipher its evolution.
The purpose of this paper is to review the Mesozoic-Cenozoic
tectonic
evolution of the Turkish erogenic belt in the framework of the theory of
plate tectonics. We discuss the ~plications of .our proposed model for the
tectonics of Tethys in the general area of the Mediterranean and the Middle
East and also for theoretical models of collisional orogeny in general. The
Tethys. “Tethys”, in general, or “Tethyan domain” refer to the large, triangular, westward-narrowing embayment of Pangaea that separated Laurasia from Gondwana-Land
and that include both Palaeo- and Neo-Tethys and their margins. As the Greek suffix
“-ide” implies. descendence, it would have been perhaps more appropriate to speak of
“Tethysides” instead of “Alpides,” but we continue to use the latter term for obvious
historical reasons. Our usage, however, deviates from the employment of Stille (1924)
and Kober (1921) in that it is geographical as opposed to theirs, which was temporal.
18-1
model is based on published literature, unpublished data made available to us
by numerous colleagues both inside and outside Turkey, and our field observations. All unreferenced
data in the following pages represent the results of
our own field work. So as not to burden the text, we have stored much of
the factual information
in Figs. 2-5 and cited sources in their captions. A
part of this information
plus some other data have also been shown in Figs.
6A-61 with the main references given in the captions.
SUTURES
AND MI~RO~ONTINENTAL
EVOLUTION
OF TIJRKEY
FRAGMENTS
INVOLVED
IN THE TETHYAN
The mountain belts of Turkey are the result of repeated continental collisions that eventually welded together the Old World segments of the two
mega~ontinents
of Laurasia and Gondwana-Land.
Since its birth, contemporaneous with the assembly of Pangaea during the late Palaeozoic, the internal
geometry of the Tethyan domain has been characterized
by a complex array
of plate boundary systems composed of a continuously
evolving network of
ridges, transforms
and subduction
zones whose record of activity is now
found, in various states of preservation,
mainly along the sutures of the
Alpides, the sites of former Tethyan oceans. The first step in reconstrL~cting
the past configurations
of continents
and oceans is therefore to identify the
sutures, which mark the places where oceans disappeared, and the continental fragments once separated by such oceans.
Figure 1 is a tectonic map showing the continental
fragments involved in
the Tethyan evolution of Turkey and their original geographic affiliations
with respect to the Neo-Tethyan
p~aeogeo~aphy.
From north to south
these are the Rhodope-Pontide
Fragment, the Sakarya Continent, the Anatolide-Tauride
Platform
and its elongated appendage, the Bitlis-Poturge
massifs. The nature of the northern boundary of the Rhodope-Pontide
fragment is controversial.
In the west, the boundary is marked largely by
deformed
TithoniaIl-~erriasian
flysch deposits (H&i et al.. 1977; Sandulescu, 1978a,
1978b)
and wild~ysch-~ontainillg
mafic volcanic
detritus
(Burchfiel,
1980) between the Rhodope Massif and the Moesian Platform.
Dewey et al. (1973) and Burchfiel (1980) have interpreted the mid-cretaceous boundary between the Rhodope Massif and the Moesian Platform as a
Fig. 1. Tectonic
setting of Turkey within the larger framework
of the eastern Mediterranean Alpides. The map was drawn from a Neo-Tethyan
perspective
(i.e., it shows only
the sutures that represent Neo-Tethyan
oceans) only to avoid too complicated
a picture.
Although the North Caspian Depression is very likely to have a ?late Devonian oceanic
basement (Fedynsky
et al., 1972; Garetsky et al., 1972; Burke, 1977; Kidd, in prep.), it
because that has been how it has
has been included in the “late Palaeozoic
platforms”
behaved since that time with respect to the Tethyan chains. A northeast striking ophiolitic
suture, the “Intra-Pannonian
Belt” of Channel et al. (1979),
may cut the basement of the
Pannonian Basin into two segments and join the Vardar Zone along the latter’s northwest
extension.
The existence
of this suture is suspected
mostly from bore-hole
data, and its
geometry
is very poorly
known.
We have left it out in our figure because its existence
could also be explained
by slicing off pieces of the northernmost
extremity
of the Vardar
Zone along strike-slip
faults that shoved a considerable
piece of the basement
of the
Dinaro3auric
Platform
into the Carpathian
bend during the Miocene
(e.g. Burchfiel,
1980).
Cimmerides
represent
the erogenic belt(s) that resulted from the elimination
of PalaeoTethys and her dependencies
(marginal basins, etc.) during the late Triassic-middle
Jurassic interval.
Only those areas in this figure have been included
in the Cimmerides,
where
the outlines
of the present
structure
were determined
by Cimmerian
deformations.
Regions
where overprinting
by later events has determined
the outlines
of the present
structures
have been left out of the Cimmeride
domain such as the Rhodope-Pontide
Fragment,
despite the fact that it contains
the main Palaeo-Tethyan
suture in this area
(Sengor
et al., in press). In spite of the current
strong deformation
in Central Iran (e.g.
Sengiir and Kidd, 1979),
it has been left within the Cimmerides,
because there the main
erogenic structures
are of Cimmerian
origin (e.g. Stiicklin,
1977). SC is Sakarya Continent
(largely
equivalent
to Brinkmann’s
(1966)
Mysisch-Galatische
Scholle),
ACB is AdanaCilicia Basin, EAAC is Eastern Anatolian
Accretionary
Complex,
R and K are Riou and
Khoura depressions
respectively.
The small circle pattern in them indicates the Oligocene
to Present molasse fill of these depressions.
Heavy lines with black triangles are subduction zones with triangles on the upper plate. Lines with half-arrows
are transform
faults
(not all of them are shown!)
and lines with hachures
on them are normal faults with
hachures
on the down-dropped
side. Short, full arrows south of the Adriatic Sea indicate
the present relative motion direction
of Africa with respect to Europe. The CarpathianBalkan regions are after Burchfiel
(1980),
whereas the Cimmeride
domains east of the
Caspian Sea are modified after Kotanski
(1978).
186
suture along which a major oceanic tract has disappeared.
Hsii et al. (1977 )
have pointed
out, however, that the “eugeosynclinal”
rocks of late Jurassic--m
early Cretaceous
age of the Balkanides
were deposited
in an extensional
basin, the Nish-Trojan
flysch trough (Hsii et al., 1977), which had originated
during the Jurassic and was most likely of modest dimensions
throughout
its
life span. The similarity
between
the Bulgarian
and the “Alpine”
Triassic
facies, the absence
of ophiolites
and oceanic deep-sea sediment,s
of Triassic
to middle
Jurassic
age in the Nish-Trojan
trough and the observation
that
the main Palaeo-Tethyan
suture is probably
located
within the circum-Rhodope erogenic
zone of Kockel et al. (1971)
and Kauffmann
et al. (1976)
(Roeder,
1978; Sengiir, 1979; Sengor et al., in press), i.e. south of the Nish-Trojan trough,
lead us to follow Hsii et al.% (1977) interpretation
as to the
nature
of the boundary
between
the Moesian
Platform
and the Rhodope
Massif, which considers
it as a narrow, intracontinental
extensional
basin prrmaturely
shut before attaining
oceanic characters
(a “sialic ocean?”
see Sen
giir and Monod, 1980).
In Turkey,
the waters of the Black Sea hide the northern
boundary
of the
Rhodope-Pontide
Fragment,
which reappears
in the Caucasus
as the flysch
zone of the southern
slope of the Greater Caucasus (Khain. 1975). Extension
here probably
started
during the early Jurassic as evidenced
by the development of the “Lias-Dagger
slatediabase
association”
(Khain.
1975).
Following
Cretaceous
flysch deposition,
the entire
belt has been deformed
through
several episodes
the latest of which continues.
Widespread
olistostrome development
in front of the southward-advancing
nappes characterizes
the Eocene. The basin was probably
eliminated
synchronously
with the Oligocene development
of the Khoura and Riou molasse basins. Although
the flysch
zone in the Caucasus
is probably
located to the south of the main PalaeoTethyan
suture (Sengijr et al.. in press), we do not believe that it ever was a
major ocean. Farther
east, the northern
boundary
of the Rhodope-Pontide
Fragment
passes across the Caspian Sea and is lost beneath
the younger
deposits of the Turan Platform
(Stijcklin,
1977). As we shall see below, palaeontologic data also indicate that the Rhodope-Pontide
fragment
was never far away
from Laurasia,
at least not since the Lias. The only exception
is probably
its
western
end, where palaeomagnetic
data show that the extreme curvature
of
the southern
Carpathian-Balkan
mountains
along the western
periphery
of
the Moesian
Platform
is the result
of the previously
straight
(east-west
trending)
Rhodopian
fragment
having “wrapped
around”
the former (Dr. N.
Orbay, personal
communication,
1979), as has already
been pointed
out by
Burchfiel(1980,
figs. 7 and 8) on the basis of geological
data.
The southern
boundary
of the Rhodope-Pontide
fragment
is marked
by
an ophiolitic
suture zone of largely earliest
Tertiary
age. In the west this
suture is represented
by the Vardar Zone (Fig. 1) (Dewey et al., 1973; Bernoulli
and Jenkyns,
1974; Biju-Duval
et al., 1977, 1978; Channel
et al.,
1979). Kockel et al. (1971) have traced the internal
limit of the Vardar Zone
in the Aegean area from Thessaloniki
through the peninsulas
of Sithonia
and
187
Athos to the island of Samothraki.
From Samothraki,
the suture strikes into
the Gallipoli peninsula (Fig. l), in the eastern part of which a late Cretaceous
ophiolitic melange crops out from beneath deformed Eocene flysch deposits
(Fig, 2) (M.T.A.,
1964 and our observations).
On the eastern shore of the
Sea of Marmara, along strike from the Gallipoli melange, Akartuna (1968)
has mapped what we interpret to be an extensive late Cretaceous ophiolitic
melange with a predominantly
pelitic matrix, on the Armutlu Peninsula. This
same melange includes blueschist slivers in the Samanli Mountains north of
Geyve, and in the Mudurnu Valley strikes into a relatively more intact, but
still internally imbricated ophiolitic complex that comprises, from bottom to
top (as restored) pyroxenites,
layered gabbros, isotropic gabbros and finally
pillow lava-chert
intercalations.
Following
a minor shale horizon a late
Cretaceous flysch was deposited on top and ?in front of the moving ophiolite
nappe. Today, the ophiolites and the flysch are found tectonically
interleaved and overthrust from the north by the so-called “Palaeozoic
of Istanbul” (Fig. 2), which here makes up the Hercynian basement of the RhodopePontide Fragment and shows a remarkable similarity to the Kraistides of Bulgaria in its stratigraphic
development
(Sandulescu,
197813). Farther to the
east-northeast,
this ophiolitic zone has been followed through Bolu and Eskipazar up to the Ilgaz Massif (Fig. 2) by Brinkmann (1966) who described the
lithologic
characteristics
of this zone as follows: “ . . . die jungmesozoischalttertiaren
Sedimente
[ sind] grossenteils als Radiolarit oder Flysch, oft mit
starker Beteiligung basischer Ergussgesteine und Tuffe ausgebildet” (p. 614).
Tokay (1973), who mapped the segment between Gerede and Ilgaz, described
a late Cretaceous
to Palaeocene
ophiolitic
melange (the Arkotdag Formation) and ascribed its origin to both sedimentary
and tectonic events that
took place in a trench associated
with a north-dipping
subduction
zone
beneath the Rhodope-Pontide
fragment. Between Gallipoli and the Ilgaz
Massif, the ophiolitic suture, herein called the Intra-Pontide
suture, separates
the Rhodope-Pontide
fragment
from the Sakarya
Continent
(Fig. 1).
Although strong magmatic arc volcanism of late Cretaceous age on the Rhodope-Pontide
fragment in this area suggests that a considerable
amount of
oceanic lithosphere between it and the Sakarya Continent had been removed,
the existence
on the latter of typically
European faunas, including such
forms as Epidocerus, indicate that this ocean was not very large.
South and southeast of the Ilgaz Massif the ophiolitic suture zone continues to delimit the Rhodope-Pontide
Fragment to the south (Fig. 1). East
of Erzincan (Fig. 2) the mainly volcanic Neogene cover hides many details of
the suture and obscures its relations to the Zangezur suture of the Lesser
Caucasus (Knipper, 1979) and also to the predominantly
oceanic rocks that
make up a large portion of the basement of eastern Turkey such as the
Yiiksekova
Complex (Perincek,
1979) and its equivalents
(Figs. 2 and 4).
East of Ankara, the Sakarya Continent wedges out and the Ilgaz-Erzincan
ophiolitic
suture zone juxtaposes
the Rhodope-Pontide
Fragment and the
Anatolide-Tauride
Platform directly. The Izmir-Ankara
Zone (Brinkmann,
188
1966, 1972, 1976) forms the boundary between the Sakarya Continent and
the Anatolide-‘I’auride
Platform (Fig. 1) and represents the remnants of a
Jurassic (?Triassic)
to early Palaeocene
ocean that closed along a northdipping subduction
zone during the late Palaeocene-early
Eocene (Din-r,
1975; Channel et al., 1979; Sengor, 197913). Because it is largely submerged
beneath the waters of the Aegean Sea, the western connection
between the
Intra-Pontide
suture and the Izmir-Ankara
zone is obscure. As the internal
boundary of the Vardar Zone strikes east-northeast
to delimit the RhodopePontide Fragment to the south in the northern Aegean, whereas its external
boundary seems to join that of the Izmir-Ankara
Zone, it appears that both
the Intra-Pontide
and the Izmir-Ankara
sutures link up with the Vardar
Zone in the west and that the Sakarya Continent was an island in the northern branch of Neo-Tethys represented
by the Vardar, Intra-Pontide,
IzmirAnkara and the Ilgaz-Erzincan
ophiolitic belts. Because the late Mesozoic--Cenozoic structural trends take a sharp bend from a predominantly
east--west orientation
to an almost northsouth
orientation at the western end of
the Sakarya Continent (e.g. Jacobshagen
and Skala, 1977), and because the
ophiolites in the island of Lesbos have this general strike, we delimited the
Sakarya Continent to the west as shown in Figs. 1 and 6B-E.
However. it is
entirely possible that it represents a somewhat enlarged eastern continuation
of the Paikon “Ridge” of the Vardar Zone (Mercier et al., 1975), with which
it shares a common tectonic position, although Eo-Hellenic ophiolite obduction (Jacobshagen
et al., 1976) is not known from the Sakarya Continent. It
is interesting
to note, on the other hand, that its predominantly
neritic:
carbonate sedimentation
of early late Jurassic age switches rather abruptly to
a pelagic environment (Fig. 5, columns 3 and 4; Fig. 6D) synchronously
with
the emplacement
of the Eo-Hellenic ophiolites onto the Paikon Ridge farther
west along strike.
The Anatolide-Tauride
Platform is the eastern end of the Dinaro-‘I’auric
Plaform as shown by the remarkable correlations
between the Hellenides and
the Anatolides and the Taurides (ijzgtil and Arpat, 1973; Bernoulli and Jenkyns, 1974; Bernoulli et al., 1974; Diirr, 1975; Brunn et al.. 1976; Diirr et
al., 1978; Channel et al., 1979; Gutnic et al., 1979; Sengiir, 1979a, b; Monod,
in press). It wedges out in eastern Anatolia in two main digitations (Fig. 1).
the Munzur Mountains in the north (ozgiil, 1976; ijzgiil et al., 1978) and the
Bitlis-Potiirge
crystalline massifs in the south (Yilmaz, 1978) (Fig. 2). These
two digitations are separated by an ophiolitic suture zone that extends from
north of the Bolkar Mountains in the southwest to Erzincan in the northeast
(Fig. 2) that contains mainly early Cretaceous ophiolites, late Cretaceous to
middle Eocene flysch and olistostromes
and that is unconformably
overlain
by latest medial Eocene sediments (Demirtasli
et al., 1973). This belt also
contains blueschists
(Metamorphic
Map of Europe, l/2,500,000.
sheet 15)
and corresponds to the eastern part of Demirtaqli’s (1977) Inner Tauride Belt.
To the southeast of this ophiolitic
suture, herein called the inner Tauride
suture, are the Malatya-Keban
metamorphics
(Periqek,
1979), which are
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ig. 2. Simplified
geological nap of Turkey emphasizing
the original palaeo-geographic
setting of the various units and de-emphasizing
me “Alpine”
metamorphism,
which, by forming extensive crystalline terrains very much confused the original stratigraphic-structural
relationships and for many years misled geological research zn Turkey a8 exemplified
by the “Zwischengebirge”
interpretation of the
An&Aides.
In conjunction
with our own observations, the map is compiled and somewhat reinterpreted from Anonymous
(1979).
Bergougnan (1976),
LX& (1975),
Geological
Map of Cyprus, l/250,000
(1979),
Geological
Map of Turkey,
l/500,000
(1961-1964),
Gonnard
et al. (1974),
Gutnic et al. (19’79),
Lapierre (1978),
Metamorphic
Map of Europe,
l/2,500,000
(1973),
0zgi.U (1976).
Perininformation
kindly provided
by J.M.N. Dyer, F. Saroglu, and Dr. N. YalGin.
gek (1979),
Tekeli (in prep.) and th e unpublished
ID = Istiranca Daglari, KaM = Kazdag
Massif, DM = Daday Massif, Ilwr = Ilgaz Massif, TM = Tokat Massif, MM = Menderes
Massif,
SD = Sultan Daglari, KM = Kiqehir
Massif, NM = Nidge Massif, MD = hlunzur
Daglari, LN = Lycian
Nappes, AN = Antalya
Nappes,
B--HN = Beysehir-Hoyran
Nappes,
H = Hadim Napges,
AM = Alanya
Massif, BD = Bolkar
Daglari, Mim = Misis Mountains,
M&f =
Malatya-Keban
Metamotphicu,
PiM = Piitiirge Massif, BM - Bitlis Massif; LE = Lake Egridir, LB = Lake Beyqehir, LT = Salt Lake (Tuz
G6lii) and LV = Lake Van.
-
czl
md
Gypsiferous
Neogme-Quaternary
mMw”o
UOlgso-Miocrnc
Noo Tethyan
Ophioktts
EASTERN
MEDITERRANEAN
ANATOLlDE/
COMPLEX
Vozgot-Sims
J
0
&tt
Km
200
/
!
AUTOCHTH.
I+---VAURIRES
LVCIAN
NAPPES
MALATYA
TAURIDES
/
/
MENRERES
----%-ANATO~I5E
METAMORWCS
MASSIF
,
--*.___
’
;
I
OY
FRAGMENT
0 N T I D ES-,
RHO0O~-~NTiO~
FRAGMENT
: RHODOPE-WNTIOE
---+j;l,c--P
‘/
SAKARYA
CONTINENT
S-----%F----PONTIDES-,
ANKARA
ZONE
imiR -
axe drawn to emphasize
the fundamentally
different
tectonic seyles of the collisional
orogen in Turkey along these crws sections (discussion in text). The Sections we drawn according
to our interpretation
of the data reported in the references cited in the caption of
Fig. 1, our observations, and Bas (1979). Canitez (1962) and Izdar (penonal communication,
1977). Note the tectonic window paFi*
tian of the metamorphio
Menderes Massif and its total disappearance
eastwards. D stands for mount of mountains. KS L the Karakaya
suture. PTS shows the probable location
of the Palaeo-Tethyan
suture along this croswwetion.
PM is khePGtiirge Maasif, EAT is thrJ
East Anatolian
tran&xm
fault, NAT is the North A~~atolian tramform
fault. Increasing ductility
down the dip of the tr;tnsform% ia
pureiy hypothetical,
but ~wne of tha miaor,
recent. crust-derived,
c&-slkalic
velcatrics
between
Erzincan and Regadiye afong the
Nurth Anatolian
transform (Bw, 1979) may be viewed a8 prndtmts of parfial melting alang the fault.
A
pp.191-192
/
1
j
1OOKm
---JO
193
composed
of metamorphosed
Permo-Mesozoic
carbonates
very similar to
those that make up the outer envelope of the Bitlis Massif (Yilmaz, 1978).
The Malatya-Keban
metamorphics
are thrust southwards over the oceanic
and island-arc lithologies of the Yiiksekova Complex, which in turn overthrust the tectonically
imbricated lithologies of the Ergani Complex and the
Poturge Massif (Hempton and Savci, in prep.) (Figs. 2 and 3B). Although the
detailed relationships
in this general area are not yet fully worked out, we
believe that the Inner Tauride suture was the original site of much of the
Yiiksekova oceanic lithologies and the combined Malatya-Keban
Pijtiirge
and Bitlis lithologies lay originally to the south of the Inner Tauride ocean.
The present distribution
of rocks is probably due to the late Eocene and
later imbrication
and thrust-stacking.
Tectonically
interleaved within and
beneath the Bitlis-Poturge
complex are late Maastrichtian to Oligo-Miocene
sediments (the Ergani Complex) that record the distensive development of a
deep-sea basin system, the Maden Marginal Sea (Fig. 6E), that disrupted the
Bitlis-Pijtiirge
Platform. In the north the Maden Basin sensu strict0 achieved
its maximum development
during the early and medial Eocene, when the
widespread
pillow lavas, radiolarian
cherts and turbidite
deposits of the
Maden Complex (Perincek, 1979; Maden Member of Rigo de Righi and Cortesini, 1964; Sason-Baykan
Group of ijzkaya, 1974; Baykan Complex of
Sungurlu, 1974 and Yalcin, 1977), and was tectonized by late Eocene times
(Fig. 5, column 16; Fig. 6F). The southern
portion,
the Ciingilg Basin,
however, remained open and closed much later, during the medial Miocene
(Figs. 6F-H).
Prior to the opening of these Maden marginal sea complexes (Figs. 6E, F),
the southern boundary between the Anatolide-Tauride
Platform and AfroArabia had been the southern branch of Neo-Tethys (Eastern Mediterranean
represents a
ocean of Sengor, 1979a). The present eastern Mediterranean
remnant of this ocean. In southeastern
Turkey and northwestern
Syria the
large ophiolite nappes of Kizildag and Baer-Bassit (Ricou, 1971; DelauneMayere et al., 1977; Yalcin, 1979), Kocali (Sungurlu, 1972; Perincek, 1979; it
would correspond
partly to Rigo de Righi and Cortesini’s (1964) Kevan
nappe), and Cilo (0. Sungurlu,
personal communication,
1979, and Y.
Yilmaz’s own field observations)
represent relicts of the closed section of
this ocean. These ophiolites were abducted onto the continental
margin of
the northern Arabian Platform as a result of its collision with the BitlisP&.&e Massifs during the late Cretaceous. Almost immediately
after this
collision event the Maden marginal sea complexes began opening.
Throughout
the evolution of Neo-Tethys
in the general area of Turkey,
the microcontinents
constantly
changed shape and position as a result of
a semicontinuous
series of rifting, shearing and collision phenomena.
The
evolution of the Antalya Nappes and the Alanya Massif (Figs. 2, 6B-E) is
a small scale example of the same phenomenon.
This makes it difficult to
define a number of microcontinents
and reconstruct their tectonic evolution
through the time period we are interested in, because such microcontinents
194
have not survived that long as individuals - ideally, different microcontinent
terminology
should have been employed for different times. This becomes
particularly
clear when we extend back into the early Mesozoic to include
the Palaeo-Tethyan
elements in reconstructions.
In Fig. 1, the Palaeo-Tethyan
sutures have not been included to keep the map reasonably simple. They are,
however, shown in Figs. 6B and C.
Finally, the eastern Anatolian
region in general represents a kind of
“suture knot” where the northern and southern Neo-Tethyan,
the Zangezur,
and the Zagros sutures come together. Until recently, there was very little
information
available on this part of Turkey. During the last several years,
however, considerable amounts of data have accumulated
and we have summarized the critical stratigraphic
information
in our Fig. 4. From this correlation chart, it appears that the basement of eastern Anatolia is largely
composed of late Cretaceous and ?older ophiolitic material. Data on pre-late
Cretaceous oceanic lithologies are now being gathered (0. Monod, personal
communication,
1980) and indicate that the eastern Anatolian oceanic area
probably originated sometime during the Jurassic (?older) which is in accordance with the data from the suture zones that converge into it. Ever since
Arni (1939) it has been known that the structure of eastern Anatolia consists generally of stacked, steepened,
imbricate thrust sheets of coloured
melange (Yiiksekova Complex and equivalents in Fig. 2) that includes, other
than Mesozoic oceanic rocks, Permian neritic limestone blocks and metamorphic knockers (see Ketin, 1977, for an excellent description
of this
association). The post-Maastrichtian
lithologies of eastern Anatolia comprise
an early flysch-olistostrome
association of Palaeocene-Eocene
age which
leads upward to less complete shallow water and finally terrestrial sedimentary sections and locally talc-alkalic volcanics. Intensity of deformation
in
these sequences decreases with decreasing age. They are often infolded or
imbricated
into the underlying
melange slices which have generally steep
contacts. We interpret the entire basement of eastern Anatolia as an accretionary wedge with at least one ?ensimatic island arc complex (andesitic
volcanism shown in Fig. 4, columns 1 and 2; these correspond to the eastern
prolongation
of the Yozgat-Sivas
Belt of Fig. 2, but are not shown there
because of the chosen scale) caught up within it. The sediments above the
melange wedge are essentially arc-trench gap and upper-slope basin deposits
recording the progressive thickening and progradation
of the wedge similar
to the Aleutian accretionary
prism with its multiple arc-trench gap basins
(Dickinson and Seely, 1979, fig. 12) or the largely subaerial Makran accretionary wedge with its large, subaerial forearc basin, the Jaz Murian depression, and the multiple, steep-sided upper-slope basins (Farhoudi and Karig,
1977, figs. 1 and 2). It is interesting to note that Arni (1939) had already
pointed out the great similarity between the lithologic associations
and
tectonic style of eastern Anatolia and those of Makran and included both of
these areas into his Iranides. Arni’s (1939) cross-sections (plate 4) across his
“geschuppte Einheiten mesozoischer und tertiaerer Schichten der Iraniden”
\
1
c ___-.-.
-
\
,
196
have an amazing similarity to those of the rotated portions of modern large
accretionary
complexes
(compare them with the sections in Karig, 19’74;
Mascle et al., 1977; Farhoudi and Karig, 1977; Dickinson and Seely, 1979).
Fig. 5. Stratigraphic
correlation
chart showing the typical successions
encountered
in the
autochthonous
and parautochthonous
regions of Turkey plus the allocbthons
of southeastern Turkey.
On the right is a summary
of major tectonic
events in Turkey
plotted
against the evidence
from those areas represented
by the columns.
The inset shows the
approximate
locations
of the stratigraphic
columns within the context
of Ketin’s (1966)
tectonic
subdivisions
of Turkey (P = Pontides,
.4 = Anatolides,
T = Taurides,
I< = Border
foids), although
they have been constructed
so as to reflect the overall traits of the sub
division in which they are located. The references
for individual columns are as follows:
1 - Ozdemir et al. (1973),
Brinkmann
(1976),
Abdiisselamoglu
(1977)
and our observations.
2 - Abdiisselamoglu
(1959)
and unpublished
observations
by Drs. A.&l. (&Zibol
and Y.
Yilmaz.
3 - Altinli (1973a),
Saner (1978)
and unpublished
observations
by Drs. A.M. (&ziibot
and Y. Yilmaz.
4 -Altinli
(1973a,
b), Bingo1 et al. (1973),
Cogulu et al., (1965),
Erk (1942)
and Saner
(1978).
5 and 6 -Agrali
et al. (1966),
Baykal (1952),
Bergougnan
(1976),
Brink~lann
(1976),
Erguvanli
(1950),
Gattinger
(1956),
Gedikoglu
(1978),
Ketin (1951),
M.T.A.
(1977)
Nebert (1961,1963),
Ozsayar (1973)
and Sengijr et al. (in press). Heavy black lines in the
Middle Jurassic.section
represent
coal measures.
7 - Boray (in Ozgiil, 1976), Diirr (1975),
Gutnic et al. (1979).
8 - Gutnic et al. (1979).
9 - Mpnod (1979).
10 - Ozgiil(l976).
11 -Argyriadis
(1974)
and Gzgiil(l976).
12 - Ozgiil(l976).
13 - Unpublished
Turkish
Petroleum
Co. section
by A. Azia, M. Meshur and H.S. Serdar
(0. Sungurlu,
written communication,
1979).
I4 - Bergougnan
(1975) and Ozgiil et al. (1978).
15 - Boray (19’75),
Yilmaz
(1978)
and Y. Yilmaz’s
observations.
The pattermess
slice
and the slice with horizontal
wavy lines beneath
the ophiolites
represent
metamorphic
aureole and the cataclastic
rocks respectively.
16 -Perinqek
(1979)
and personal communications
by Dr. D. Perirqek
and 0. Yungurlu
(1979).
17 - Rigo de Righi and Cortesini
(1964),
Sungurlu
(1974)
and written communization
by 0. Sungurlu (1979).
Key to legend:
I = deformed and metamorphosed
stratigraphic
basement,
2 = the “Palaeozoic of Istanbul,”
3 = late Palaeozoic-early
Mesozoic
granitic intrusives,
3 = conglomerates and breccias of various kinds, 5 = olistostromes,
6 = sandstones
of unspecified
sedimentologic
environment,
7 = flysch, 8 = shales, 9 = pelites and semi-pelites
of the Alanya
and Bitlis Massifs, 10 = evaporites,
1 I = lagoonal and limnic carbonates,
12 = neritic lime15 = radiolarites,
16 = magmatic
arc volstones,
13 = dolomites,
14 = pelagic limestone,
canism,
17 = magmatic
arc plutonism,
18 = rift volcanism
(mixed rift and arc volcanism
symbols indicate bimodal, Tibetan volcanism
in column 6), 19 = pillowed basalts, 20 = op23 = no record. S represents
serhiolites,
21 = mdlange, in part ophiolitic,
2, 7 = bauxites,
pentinite
in the columns;
small, open triangles, chert nodules or cherty beds. LN = Lycian
Nappes, B-HN
= Beygehir-Hoyran
Nappes,
HN = Hadim Nappes.
Figures give approximate thicknesses.
5
g
UPPER
e MIDDLE
s
LOWER
UPPER
ii!
w MIDDLE
u
LOWER
3
i!!
G
g
v
g
UPPER
I
LOWER
UPPER
2 MIDDLE
oc
,3 LOWER
UPPER
%
$
MIDDLE
F
LOWER
6
7
8
15
a,
P
BLACK
EASTERN
SEA
MEDITERRANEAN
.,’
:
.‘.
‘”
. .
:.
;. ‘: : 1
.- i: ‘v
L
:
‘“‘..
..:.
,111
.
16
17
pp. 197-202
vv v
16 vv
vvv
L-l
latdialArabia
c.ollisicn
al stackii
of the Anatdiil
Tarkle Platform and the
generation of
the Andoliis
ion of Pontide; with
AmtoMe/ Tauride
Platform
the
#%olite abduction onto the
Anatoliil
Tour’& Platform
ure of Palaeo-Tethys
The’Lissic
Event’ in the Tauides
-Opening of the Northern &onch
of Neo-Tethys
hosue
of Karr~kaya Margiml
Sea
Wpening of the BiWZqms/,
Eastern Mcdtcrmncan Ocean
(Southern Branch of Neo-TethysI
Openirq of Komkaya Hardnal Sea
203
In this paper we call the accretionary
prism or prism system of eastern Anatolia the East Anatolian Accretionary
Complex following Sengor et al. (in
press) (Figs. 1 and 4). As we shall see below, the East anatolian Accretionary
Complex exercised a very profound influence on the neotectonics
of Turkey
(Sengor and Kidd, 1979; Sengor, 1980).
As to the positions of the various continental fragments at different times,
little can be said for the time being. Existing palaeobiogeographical
and
palaeomagnetic
data on Turkey in general allow us only to make the very
crudest estimate concerning the sizes of oceans and locations of continental
fragments.
The Rhodope-Pontide
Fragment and the Sakarya Continent
share Eurasian faunas since at least the Lias and contrast sharply from the
Lias to the early Tertiary with the Anatolide-Tauride
Platform and Arabia,
which display southern Tethyan faunas during this time interval (e.g. Bassoullet et al., 1975; Bergougnan, 1975; Fourquin, 1975; Enay, 1976; Gutnic
et al., 1979). This probably indicates that the ocean today represented
by
the Izmir-Ankara-Ilgaz-Erzincan
ophiolitic suture was fairly wide. That
the Sakarya Continent was never very far away from Eurasia is also supported
by preliminary palaeomagnetic
data on its palaeolatitudes
during the Jurassic (I. Evans, personal communication,
1980). However, it must have been at
least a few hundred kilometres away from the Rhodope-Pontide
Fragment,
because the elimination
of the intervening ocean by northdipping
subduction beneath the Rhodope-Pontide
Fragment during the late Cretaceous
created a well-developed island-arc on top of it.
Neither the faunas (Bassoullet et al., 1975) nor, it appears, palaeomagnetics
(Zijderveld and Van der Voo, 1973; Channel et al., 1979) are aware of the
existence of the southern branch of Neo-Tethys.
This is understandable,
because it terminated
westwards around Sicily, not far away from Turkey
and, as the palaeomagnetic
data do not show its existence, its angular opening must not have exceeded
the inherent 10” error in the data, which
amounts to about a 700-km north-south
opening between Arabia and eastern Anatolia. After the late Cretaceous collision between the Bitlis-Potiirge
Massifs and the Arabian Platform, faunal exchange between the latter and
the Anatolide-Tauride
Platform must have been facilitated.
Palaeomagnetic
evidence for the successive positions of the Cimmerian
Continent in Turkey since the late Palaeozoic is extremely scanty, and palaeobiogeographic data are nonexistent.
Zijderveld and Van der Voo (1973) concluded that during the Permian the entire present Turkish area belonged to
Africa, a conclusion entirely consistent with the geological evidence.
In the following paragraphs we outline the tectonic evolution of Turkey in
terms of nine selected time intervals. These intervals have been selected to
show the major changes during the tectonic evolution of the country and do
not necessarily reflect data resolution. With the available data from Turkey
one could draw much more detailed palaeogeographic
and palaeotectonic
maps for much shorter time intervals. Such an attempt would be far beyond
the scope of this paper, whose purpose is simply to present the outlines of
204
the geological development in terms of plate tectonics.
In the palaeotectonic
maps the present shapes of the microcontinental
blocks have been dotted, following Burchfiel’s practice (1980), for reference
purposes. Very large amounts of internal strain characterize all of the microcontinental
blocks shown in our Fig. 6 and that is why we avoid the term
“plate” for such blocks. Large strains generally begin, however, only after
continental
collision has brought together two continental
pieces along a
suture zone. As long as a continental
block is located in a plate with only
oceanic boundaries,
little distortion
occurs in it. In our palaeotectonic
maps we did not indicate the positions of ridges. There are very few data
available to do this and unless good transform fault strike control exists (as
in Fig. 6A, for example) to constrain spreading directions, they are nearly
impossible to use for making unique reconstructions
of ridge geometries.
The original shapes of the microcontinents
shown in the palaeotectonic
maps of Fig. 6 are poorly constrained. We assumed, largely for graphic reasons, an overconservative
40% total across-strike shortening in Turkey since
the Permian (excluding the width of subducted oceans) and extended the
respective pieces in a northsouth
direction to restore back that amount of
area. The real amount of shortening in Turkey since the Permian, when the
amount due to oceanic subduction is taken out, is probably closer to about
60% or more. The amount of post-Miocene east-west
shortening that has
affected Turkey (Zjengiir, 1979b, 1980) has not been taken into account at
all and left unrestored.
As long as oceanic lithosphere is present between the
continents the distance by which they are apart has been treated, for practical purposes, as unknown and the oceans in our Fig. 6 are drawn only symbolically. Only in the case of the Neo-Tethyan
oceans have we roughly
attempted to show relative size.
PERMO-TRIASSIC
EVENTS
(Fig. 6A)
During the Permian, the entire area of present Turkey constituted a part of
the northern margin of Gondwana-Land
facing Palaeo-Tethys. $engor et al. (in
prep.) recently summarized the available data on the nature of this margin.
In the eastern part of the eastern Pontides *, in the G~m~~hane-~ayburt
area a thick marine sequence (approximately
1.5 km) composed
of red
arkoses, orthoquartzites,
and fossiliferous dark limestones with interlayered
hornblende-biotite
andesites, tuffs and other mainly silicic lavas was deposited during the Permo-?Carboniferous
over a metamorphic
basement. This
basement consists of polyphasedeformed
quartz-feldspathic
schists, phyllites, and slates (Yilmaz, 1974a). Post-kinematic,
low-melting composition
granites, such as the Giimiishane granite pluton, intruded this basement.
Whole rock-Pb ages on this pluton range from 298-338
m.y. Sengor et al. (in
* For Turkey’s
tectonic
Taurides and the Border
subdivisions
as defined by Ketin
Folds, see the inset in Fig. 3.
(19661,
Pontides,
Anatoiides,
205
press interpreted
the Permian andesitic volcanism as products of a southdipping subduction
zone that was consuming Palaeo-Tethyan
ocean floor.
Whether the intrusion of the Giimiighane pluton is also related to this or to
some older “Hercyanian”
regime is as yet uncertain. If the age of the flyschsequences overlying Palaeo-Tethyan
ophiolites in the Kiire, Cangal Dagi and
Daday regions (Sengijr et al., in press) extends down into the Permian as
hypothesized
by Ketin (in M.T.A., 1962), they may be taken as evidence for
Permian flysch deposition
on Palaeo-Tethyan
oceanic crust. In this region
there is documented Triassic flysch. Associated with the tectonized ophiolites
are blueschists and eclogites that are of pre-Dogger age. In the rest of the eastem Pontides, Triassic rocks are very sparse, but because of the apparent continuity between the Permian and Jurassic tectonic events, Qengiir et al. (in
press) assumed that there had been no significant change in the tectonic
regime of the eastern Pontides during the Triassic.
In the northern part of the western Pontides, the Permian saw the end of
the late Palaeozoic Hercynian deformations.
The coastal regions of the Black
Sea were the site of continental
red bed deposition
(Brinkmann,
1976),
and were probably located at some distance to the south of the active
Palaeo-Tethyan
margin. With the exception of the Orhanlar Greywacke, a
late Carboniferous-early
Permian marine molasse related to the waning Hercynian deformations
and located to the south of the Sea of Marmara (Brinkmann, 1971, 1976), the “porphyroids
and the Verrucano of Sandikli” (Parejas, 1943a; Gutnic et al., 19’79), and the predom~antly
elastic Lower Permian section of southeastern
Turkey (Rigo de Righi and Cortesini, 1964),
the rest of the country was the site of neritic carbonate deposition indicating
quiet platform conditions.
During the Triassic this platform ruptured to produce extensional basins
in two areas at two different times. During the early Triassic a rifting event
began along a line that extends from the Biga Peninsula north of Bursa
through Bilecik and Ankara to the Tokat Massif (Bingiil, 1976). It may extend
farther east in the direction of Erzincan (0. Tekeli, personal communication),
but the present data there are equivocal.
On the Biga Peninsula Sing61 et al. (1973) and Bingo1 (1976) described a
series of spilitic basalts, mud&ones and radiol~ites
that interfinger with
sandstones, siltstones and conglomerates,
collectively termed the Karakaya
Formation.
The blocks within the conglomerates
contain pieces of the Permian neritic carbonates.
The Karakaya Formation
is overlain by medial
Triassic sediments (Bingo1 et al., 1973). Bingo1 (1976) interpreted
the early
Triassic Karakaya Formation as indicative of rifting of the previously extensive Permian carbonate platform and subsidence of the rift floor. He assumed
that the early Triassic rifting here never produced real oceanic crust, because
no evidence for the existence of a well-developed
ophiolite suite had been
found in the Karakaya Formation.
Tekeli (in prep.) however, has later
pointed out the existence of nearly all the members of the ophiolitic association here and thought the Karakaya basin to have been underlain by real
206
Fig. 6. Palaeotectonic
maps depicting
the tectonic
evolution
of Turkey and some neighbouring regions from the Permian to the present.
Maps show oceans (widths are not to
scale and shown only symbolically;
see text;
horizontally-ruled
= Palaeo-Tethys
and
dependencies;
vertically
ruled = Neo-Tethyan
oceans),
continents
(white),
predominant
lithologic
and lithohcies
types (widely
spaced carbonate
pattern = neritic
carbonates,
closely spaced carbonate
pattern = pelagic carbonates,
J = evaporites
of central and eastern Anatolia,
c = coal; open arrows are very generalized
sedimentary
dispersal directions
for elastics
(no pattern = Atlantic-type
continental
margin turbidites,
unless otherwise
specified,
F = flysch)
solid black circles are blueschist
facies metamorphics,
half-filled
circle is eclogites,
whereas g represents
greenschist,
an amphibolite
facies metamorphics,
R = tholeiitic
basalt, AB = alkalic basalt, B = unspecified
basalt, 7% = trachyte,
c = arc volcan&,
++ = arc plutonies,
A = Tibetan-type
volcanics,
s = untectonized
ophiolites.
Thin
lines with hachures
on them are Atlantic-type
continental
margins (normal faults in Fig,
611, heavier lines with black triangles
on them are subduction
zones (triangles
on the
upper plate) and lines with half-arrows
are transform
faults. Lines with open triangles
show major intracontinental
thrust faults (triangles on the upper plate), lines with upsidedown triangles
portray
retrocharriage
(triangles
on the lower plate), lines with doubleheaded arrows across them are folds, and the normal fault + conglomerate
symbol indicates rifting.
Ladder-patterned
solid lines are sutures.
The main references
to the data
used for individual time frames are as follows (discussion
in text):
A - (Permo-Triassic).
Bingo1 et al. (1973),
Bingo1 (1976),
Brinkman
(1976),
Cogulu
(1975),
Cogulu et al. (1965),
D&r (1975),
Eren (1979),
Erk (1942),
Erol (1953),
Giirpinar (1976),
Gutnic
et al. (1979),
Ketin (1951),
Lisenbee
(1971),
Marcoux
(in press),
M.T.A. (1964),
Ozgiil (1976),
Ozkocak
(1969),
Perinc;ek (1979),
Rigo de Righi and Cortesini (1964),
Robertson
and Woodcock
(1979),
Saner (1978),
Sung&u
(1974),
Van der
Kaaden (1959,
X966), Yilmaz (1972,1973.1974a,
b).
B - (Early Jurassic).
Adamia et al. (1977),
Argyriadis (1974),
Baykal(1952),
Bergougnan
(1975),
Fourquin
(1975),
Gedikoglu
(1978),
Gutnic et al (1979),
Kauffman
et al. (1976).
Ketin (19511, Monod (1979),
M.T.A. (1964),
Nebert (1963),
Ozgiilf1976),
Rigo de Righi
and Cortesini
(1964)
8 en g or et al. (in prep.), Seymen (1975),
Wedding (19631,
Yilmaz
(1972).
C -(Middle
Jurassic).
Adamia et al. (1977),
Agrali et al. (1966)
Gutnic et al. (1979),
Khain (1975),
M.T.A.,
1962 and Nebert (1961),
Ozgiil (1976),
Rigo de Righi and Cortesini (1964),Seymen(1975),pengor
et al. (in prep.) and 0. Yilmaz(1978).
(See p, 213.).
D - (Late Jurassic--early
Cretaceous)
Adamia et al. (1977),
Akin (1979),
Altinli (1973a,
b), Batman (1977),
Baykal (1952),
Brinkmann
(1972,1976),
Gutnic et al. (1979),
Ketin
(lS51),
M.T.A. (IS62),
Gzgiil(1976),
Wedding (1963)
and Zankl(lS61).
(See p, 215.).
E .- (Late Cretaceous-Palaeocene).
Adamia et al. (1977),
Al Maleh (1976),
Bergougnan
(1975),
Bingo1 (1976,
1978),
Brinkmann
(19?6),
Brunn et al. (1971),
Cogulu flS67),
~alapkulu
(1978),
Demirtasli
et al. (1973),
Diirr (1975),
Fourquin
(1975),
Gedikoglu
(1978), Gutnic et al.(lS79),
Hall (1976),
Lapierre (1975),
Letouzey
et al. (1977),
M.T.A.
(1962),
&&I
(1976)
Ozgiil et al, (1978),
Perineek
(1979f,
Rigo de Righi and Cortesini
(1964),
Saner (1977,
lS78),
Seymen
(1975),
Sungurlu
(1974),
Taner (19771,
Tekeli
(1978),
Tokel (1977),
Yilmaz (1972,
1978) and the unpublished
data both furnished by
Drs. N. C&jr%-, D. Peringek and Messrs E. Arpat, M. Hempton, N. Ozgul, G. Savci and 0.
Sungurlu and collected
by ourselves. (See p. 217.).
F - (Early-middle
Eocene).
Abdiisselamoglu
(195S),
Ataman
(1972),
Bergougnan
(1975), Cogulu (1975),
Demirtasli
et al. (1973),
Diirr (19751,
Gedikoglu
(19781,
Gutnic
et al. (1979),
Izdar (1975),
Kalafat$ioglu
and Uysal (1964),
Norman
(19731,
Ozgul
(1976),
Rigo de Righi and Cortesini
(1964),
Saner (1977),
Seymen
(19751,
Sungurlu
(1974),
Tokay (1973),
Tokel(1977),
Vachette
et al. (1968).
(Seep.
222.).
G -(Late
Eocene and early Miocene).
Bingo1 (1976),
Diirr (1975)
Gutnic et al. (1S79),
Luttig and Steffens
(lS76),
ijsgiil (1976),
Sungurlu
(1974)
and unpublished
data by Dr.
207
A
NORTHEAST
AFRICA
GONDWANA-LAND
with the
Apulian
“0
6
-
3,il
!~
ig. 6A. Palaeotectonic
/,I, j!i !I(:/j!j,l, i;j/ii ‘f
map of the Permo-Triassic. B. Palaeotectonic
_
map of the ea
D. Periqek, Messrs. M. Hempton, 0. Swngurlu and ourselves. (See p. 224.).
Ltittig and Steffens (1976), SengGr (1978,
H - (Middle Miocene-Pliocene).
1980), gengijr and Dewey (in press), gengiir and Kidd (1979) (See p. 226.).
I - (Pliocene-present).
Same as Fig. 6 H (See p. 228.).
1979b,
oceanic lithosphere.
The Upper Triassic in the Karakaya basin is in flysch
facies and indicates syntectonic
deposition.
The basin closed before the
beginning of the early Jurassic and the Lias lies unconformably
on top of the
north-vergent
structures of the deformed Karakaya basin with a basal conglomerate (Radelli, 1970; Fourquin, 1975; Bing61, 1976). The deformation
of the Karakaya basin produced a typical ophiolitic mklange involving all the
lithologies of the Karakaya Formation,
metamorphosed
under a variety of
conditions producing blueschist, greenschist and high-grade Barrovian assemblages (Tekeli, in prep.). Because of its location behind the Palaeo-Tethyan
trench and arc and its relatively short life span, we regard the Karakaya basin
as a marginal sea that opened above the Palaeo-Tethyan
subduction zone. We
have been unable to trace the Karakaya suture either t,o t,he west of the
Sakarya Continent or to the east of the Tokat Massif. This may be either
because of lack of data or the original termination of the basin, perhaps like
the present Black Sea.
Throughout
southern and southeastern
Turkey, as in many other places
along the margins of the eastern Mediterranean,
there is evidence for a
Ladinian-Norian
rifting event. In the Antalya Nappes (Delaune-Mayere
et
al., 1977; Marcoux, in press) the early Triassic has a neritic carbonate
environment
similar to the Permian platform regime. In the late Anisian,
development
of breccias with pelagic matrix and Daonellidecontaining
pelagic limestones with red, manganiferous radiolarites of Ladinian age mark
the disintegration and subsidence of the platform. Pietre Verdi (mainly intermediate-silicic
tuffs) volcanism at the Anisian-Ladinian
boundary (a very
common characteristic of the entire Dinaro-Tauric Platform) is followed by a
strong, predominantly
alkalic basaltic volcanism
of late Carnian--early
Norian age in the area, which produced abundant pillow-lavas. This abundant
late Triassic mafic volcanism is also present in the Mamonia Complex of
Cyprus (Lapierre, 1975; Lapierre and Rocci, 1976). As equivalents of the
Antalya Nappes are present beneath the allochthonous
body of the Alanya
Massif (Gutnic et al., 1979), the rifting of the latter from the Anatolide~Tauride Platform must have begun at this time as well, to open at least a part
of the basin in which the lithologies of the Antalya Nappes were deposited
Turkey evi(Basin Pamphylien
of Dumont et al., 1972). In southeastern
dence of Carnian-Norian
rifting and subsidence of the carbonate platform is
recorded in a number of places within the lowest tectonic slice of the Bitlis
Massif. To the south of Bigra Dagi, for example, submarine basaltic flows
interfinger with Megalodont-bearing
marbles and pass upwards into red siltstones and cherty limestones. Local development
of polygenic breccias here
may be indicative of active faulting. (D. Perinqek and 0. Sungurlu, personal
communication,
1979). Within the Triassic successions of the autochthon
of
the Border Folds region, i.e. the northern part of the Arabian Platform, Sungurlu (1974) has shown that sedimentation
was controlled by active normal
faulting.
Elsewhere around the eastern Mediterranean
there is other, albeit indirect
209
evidence for the Triassic rifting of Turkey away from Gondwana-Land. Friedman et al. (1971), Goldberg and Friedman (1974), and Bein and Gvirtzman
(19771, have shown that during the late early Jurassic a continental shelfslope-rise triplet existed along the Levant coast of Israel and they have
inferred that the same geometry must have existed also during the early
early Jurassic in the area. G.M. Friedman (personal communication, 1978)
believes this to be a result of a late Triassic rifting event here. Farther south,
in the Sinai area Ginzburg and Gvirtzman (in press) have shown the existence
of a similar picture. This and other complementary evidence from the western part of the eastern Mediterranean (W.B.F. Ryan, personal communication, 1979) show that the opening of the eastern Mediterranean probably
began during the Carnian-Norian interval with some preliminary events
during the latest Anisian-Ladinian (Pietre Verdi time).
Although Dewey et al. (1973) had already inferred the Triassic opening of
the eastern Mediterranean ocean, which they had incorrectly connected with
the Vardar Ocean, some later workers denied its oceanic nature. The existing
equivocal geophysical data have been interpreted by some authors to indicate
a continental-type basement beneath the present eastern Mediterranean
(Morel& 1973; Morelli et al., 1975; Lort, 1977; Woodside, 1977). On the
other hand, the Orsay group of geologists have primarily developed the concept that all the ophiolites in the eastern Mediterranean have their roots in
the Izmir-Ankara-Ilgaz-Erzincan-Zagros
ophiolitic suture and that they
have been thrust over the Anatolide-Tauride Platform during the late Senonian (Ricou et al., 1974,1975,1979;
Delaune-Mayere et al., 1977; Gutnic et
al., 1979). This inte~retation considers the eastern Medite~anean as an
intracratonic rift or a foredeep (Brunn, 1979) that never had any oceanic
substratum. Other workers following this view have argued that the NeoTethyan evolution of the eastern Mediterranean ranges has been the result of
the closure of a single Neo-Tethyan ocean to the north of the Dinaro-Tauric
Platform and that the latter has always been an integral part of Afro-Arabia
(e.g. Channel and Horvath, 1976; Laubscher and Bernoulli, 1977; Channel et
al., 1979). This argument was originally developed primarily to explain the
great similarity in the timing of ophiolite emplacement in diverse parts of the
eastern Mediterranean orogen (L.-E. Ricou, personal communication, 1978).
As a result of recent field work on the ophiolites of southeastern Turkey and
northwestern Syria, however, the age of emplacement of these ophiolites
became somewhat older (Camp~i~)
than the mainly late Senoni~ nappes
of the Anatolide-Tauride Platform (Al Maleh, 1976). However, the most
decisive evidence against a supra-Anatolide-Tauride origin for the Cypriot
and southeastern Turkish ophiolites comes from the central part of the
Anatolide-Tauride Platform itself. Near Sariz, the geologists of the Turkish
Petroleum Company recently discovered a complete stratigraphic sequence
from Upper Jurassic to Middle Eocene (Fig. 5, column 13) with no evidence
of ophiolite nappes passing over the platform during the late Cretaceous.
This automatically puts the roots of the eastern Mediterranean ophiolites
210
south of the Anatolide-Tauride
Platform and shows that even if the present
eastern Mediterranean
does have a “continental”
substratum (see Sengor and
Monod, 1980), it must have had an oceanic floor at least during the Cretaceous. This resolves the controversy
over the origin of the eastern Mediterranean. It is an ocean that began opening during the late Triassic and partially closed, along the Bitlis Suture (Sengiir et al., 1979) during the middle
Miocene (see below).
The Permo-Triassic
geologic record of Turkey records the subduction of
the Palaeo-Tethyan
ocean floor along a south-dipping
subduction
zone
beneath Turkey. This subduction gave rise to the opening of the Karakaya
marginal sea during the early Triassic, which closed shortly thereafter during
the late Triassic. The eastern Mediterranean
began opening during the Carnian-Norian
interval, an event that marks the opening of Neo-Tethys in this
area; towards the east the opening continued
through the Zagros Ocean
(Stocklin,
1974, 1977) all the way into the Himalayas, separating a northern
strip, the Cimmerian Continent
(Sengor, 1979a),
from Gondwana-Land
as
Neo-Tethys opened in its wake.
EARLY
JURASSIC
EVENTS
(Fig. 6B)
As far as the Palaeo-Tethyan
palaeogeography
is concerned,
the early
Jurassic events represent a continuation
of the Permian regime. However, at
this time the ocean floor of Palaeo-Tethys
was receiving more abundant
flysch sediments as they are today seen above the preserved Palaeo-Tethyan
ophiolites all the way from the Artvin area near the Russian border (Sengor
et al., in prep.) through the Kelkit exposures, to the Ktire-Daday region
(Ketin, in M.T.A., 1962b; Sengor et al., in press). Because the Palaeo-Tethys
terminally
closed during the middle Jurassic (Fig. 6C), the increased elastic
influx in the early Jurassic may reflect partly the presence 01 an approaching
northern
continent,
most likely the Scythian
Platform.
Although in the
Lesser Caucasus
strong early Jurassic
talc-alkalic
volcanism
was active
(Adamia et al., 1977), in the eastern Pontides it assumed a subordinate position with respect to an increasing tholeiitic volcanism (Yilmaz, 1972). Concurrent with this intense tholeiitic
activity, rifting began near the volcanic
axis of the north-facing
Palaeo-Tethyan
arc. This could perhaps be interpreted as a result of the subduction
of progressively
older oceanic lithosphere. While the Laurasian continental
margin moved closer to the magmatic arc, the latter became an extensional
arc (Dewey, 1980), along the
volcanic axis of which marginal basins began to open. Within the initial rift
trough very thick conglomerate
deposits (>l km) had intercalated tholeiitic
basalts and trachytes (Bergougnan,
1976).
During the Sinemurian a marine transgression began from the south (Akin,
1979)
onto the Palaeo-Tethyan
arc (Ketin, 1951),
which had become an
island arc due to the early Jurassic rifting, indicating the beginning establishment of a south-facing
Atlantic-type
continental
margin here. This early
211
Jurassic rifting south of the eastern Pontides is well documented all the way
from the east of the Ilgaz Massif to the northeast of Erzincan (Seymen,
1975; Bergougnan, 1976; Tokel, 1977). This is also the time when faunal
differentiation between the Pontides and the rest of Turkey began (Bassuolet
et al., 1975; Enay, 1976), which also indicates that the Palaeo-Tethys had
been reduced to a size that permitted the Laurasian faunas to reach the
Pontides. In fact, evidence from the Mashad suture in northern Iran indicates
that there the ocean had already closed by the time of the deposition of the
(?Rhaeto-Liassic) (StScklin, 1974, 1977; Majidi,
Shemshak Formation
1978). The equivalent of the Shemshak in the Pontides is the coal-bearing
Kelkit Formation of Bergougnan (1976), where it occupies a back-arc setting,
along the northern margin of the opening northern branch of Neo-Tethys.
Along the Intra-Pontide suture and the Izmir-Ankara zone the evidence
for the Liassic rifting event is less abundant than in the eastern Pontides.
South of the Mudurnu area the early Jurassic basaltic volcanics of tholeiitic
affinity are overlain by typical shelf deposits of medial to late Jurassic age
(Altinli, 1973a) and present a picture very similar to that observed in the
eastern Pontides. Fourquin (1975) points out that the Izmir-Ankara ocean
also began opening during the Lias (or earlier) as this is the time of faunal
differentiation as it was in the eastern Pontides. This widespread early Jurassic rifting event is recorded also in the Vardar Zone and in the rest of the
peri-Dinaro-Tauric Platform oceans (Channel et al., 1979) with the exception of its southeastern margin that had already rifted during the Triassic.
The opening of this new arm of Neo-Tethys, herein called the northern
branch, greatly reduced the width of the Cimmerian Continent in this area
and isolated the Dinaro-Tauric Platform, except in Sicily where there has
never been an oceanic gap between Africa and the Dinaro-Tauric Platform, as
shown by the continuity of the facies belts from Africa to Italy (e.g. Channel
and Horvath, 1976, Caire, 1977; Channel et al., 1979).
Unfortunately, we have little evidence on the Anatolide-Tauride Platform
rifting away from the Cimmerian Continent during the early Jurassic, because,
its northern margin has been heavily tectonized during the later collision and
the evidenced has been destroyed.
The similarity between the basement of the Bitlis-Poturge massifs and the
eastern Pontides and their latest Palaeozoic sedimentary covers (compare the
lowest sections in Fig. 5, columns 4 and 15) lead us to believe that the continental piece that was rifted from the eastern Pontides east of the Munzur
Mountains was the Bitlis Massif. The fact that its late Triassic-early Jurassic
cover sediments consist mainly of neritic carbonates may be because the
Palaeo-Tethyan arc at this time had become an extensional arc with mainly
mafic volcanic constructions of low relief that did not supply much elastic
material to surrounding areas. As the Bitlis Massif was rotating away from
the Pontides it ripped the Poturge and Malatya-Keban massifs and the Bolkar
Mountains away from the rest of the Anatolide-Tauride Platform. The timing of this rifting event that opened the Inner-Tauride ocean is very poorly
212
constrained
but sparse evidence of alkalic magmatism at the northeastern
end of the rift zone indicates a Jurassic opening (I. Seymen, personal communication., 1980).
Early Jurassic is a time of emergence in many parts of the Central Taurus
and has witnessed the deposition of Liassic polygenic conglomerates
(Cayir
and Uziimdere Formations: Gutnic et al., 1979; Monod, 1979) now found in
the autochthon of the eastern flank of the Gulf of Antalya. Argyriadis (1974)
has pointed out, (following a joint excursion with N.ozgtil) the simularities
between the Alanya Massif and the Bolkar Mountains with respect to their
stratigraphic
content and geological evolution. Both areas particularly
display the Liassic emergence as do the auto~hthonous
series just north of the
Alanya Massif (Fig. 5, columns 10-12). This suggests the possib~ity that the
Inner-Tauride
ocean may have connected
with the Pamphylian basin that
had already separated the Alanya Massif from the Anatolide-Tauride
Platform during the late Triassic. This rifting event may explain the rather
abrupt early Jurassic emergence in this localized area and the Cayir and
Uziimdere conglomerates
can be viewed as correlative deposits generated by
the erosion of the uplifted rift shoulders. This hypothesis (N. &@il, personal
communication,
1978) implies that during the Lias the combined AlanyaBolkar-Malatya-Keban-Poturge-Bitlis
continental
fragment began to be
separated from the main body of the Anatolide-Tauride
Platform and supports &$il’s
(1976) grouping of these units under a single heading, his
Alanya Unit. It is unfortunately
almost impossible to directly test this
hypothesis
as the critical region between the Bolkar Mountains and the
Rlanya Massif are buried beneath the Neogene fill of the central Anatolian
ovas and the Miocene marine limestones of south--central
Turkey (Fig. 2).
In summary, the early Jurassic was a time of continued disintegration
of
the Cimmerian Continent in Turkey that gave birth to the Anatolide-‘I’auride
Platform and possibly another independent
continental
fragment, that of
Alanya-Bolkar
Mountains-Malatya-Keban-Poturge-Bitlis.
At this time
the southern branch of the Neo-Tethys presumably continued its growth and
the northern
branch originated
as a Palaeo-Tethyan
marginal basin and
extended all the way into the present western Mediterranean area.
MEDIAL
JURASSIC
EVENTS
(Fig. 6C)
The most important change that occurred during the middle Jurassic in
Turkey (and in a large portion of the rest of the Tethyan domain between
the Balkans and China) was the terminal closure of Palaeo-Tethys
which
resulted in the collision of the Cimmerian Continent with the Scythian Platform (Sengijr et al., in press). In the eastern Pontides, the northern areas of
the magmatic arc emerged and sedimentation
was interrupted (Ketin, 1951;
Bergougnan, 1976) while in the southern portions it continued, albeit in a
regressive facies, without any serious interruption
(Agrali et al., 1966).
Flysch deposition on the Palaeo-Tethyan
oceanic crust ceased during the late
213
PLATFORM
Fig. 6C. Palaeotectonic
map of the middle Jurassic. For legend see p. 206.
Lias-early
Dogger and very shortly thereafter both the sediments and their
ophiolitic
substratum were overridden from the south by the continental
basement of the eastern Pontides. This event has been interpreted in terms
of the collision between the Scythian Platform and the eastern Pontide basement (a part of the Cimmerian Continent) (Sengor et al., in press). The collision gave rise to a north-vergent two-phase penetrative deformation,
the
second one being accompanied by a greenschist metamorphism in the oceanic
lithologies. The Cimmerian Continent itself was not penetratively deformed,
except for a wide zone of intense cataclasis and mylonite generation along its
basal thrust (Sengor et al., in press). The thrust contacts were later intruded
by a variety of plutons that range in composition
from granodiorite to tonalite and give isotopic ages around 165 m.y. (Yilmaz, 1979). This event also
seems to have reset the K/Ar clocks of the Giimiishane pluton, which indicate
a thermal event around 165 m.y. ago (Cogulu, 1975). Intensive north-vergent
deformation,
granite intrusion and greenschist metamorphism
of probable
214
Dagger age has obliterated
Palaeo-Tethys
along the ~ircum-Rhodope
orogen
(Kockel et al., 1971; Kaufmann et al., 19’76; Roeder, 1978) as well, where
Roeder (1978) has identified a north-vergent ophiolite abduction of this age.
With the terminal closure of Palaeo-Tethys
all oceanic and most marine conditions terminated
in the peri-Black
Sea regions until the opening of the
Black Sea itself during the Cretaceous (Brinkmann,
1974).
In the rest of Turkey carbonate
deposition
and quiet platform--shelf
development dominated the scene. Only to the south of the Munzur Mountains an as yet poorly understood
and poorly dated deformational
event
provided some variety for the otherwise rather monotonous
course of events.
This deformation
is expressed in an angular uncomformity
where lipper
Cretaceous covers folded Jurassic rocks (I. Seymen, personal communication,
1980). Although the Neo-Tethyan
oceans presumably continued to expand,
we have no way of knowing how large they were during the middle Jurassic.
LATE
JURASSIC-EARLY
CRETACEOUS
EVENTS
(Fig. 6D)
The late Jurassic and the early Cretaceous events in Turkey are intimately
connected in time, thus we treat them together.
Following the closure of Palaeo-Tethys
during the middle Jurassic, continued convergence
led to crustal thickening
and Tibetan-type
volcanism
(i.e., volcanism caused by the partial melting of the lower crust and its fracturing perpendicular
to shortening due to collision:
see Dewey and Burke,
1973) throughout
the eastern part of the eastern Pontides and the entire
Caucasus area, In the northern part of the Lesser Caucasus the talc-alkalic
volcanism terminated during the late Jurassic and was replaced by localized
alkali basalt-+rachyte
lavas intercalated
with salt- and gypsum-bearing
terrestrial deposits (collision related rifting? see !$eng6r et al., in press), which
were coeval with ongoing talc-alkalic
volcanism in the south (Adamia et al.,
1977).
A similar setting is shared by the Neocomidn bimodal volcanism of
the northern Giimiiahane area (Zankl, 1961). The abundant Bajocian-Bathonian granitic plutons that are found both in the Main Ranges of the Greater
Caucasus and in the basement of the Trans-Caucasus
depression (Khain,
1975)
may be the earliest manifestation
of the “Tibetization”
process in
the area.
Throughout
the late Jurassic--early
Cretaceous carbonate shelf deposition
shooting
and the growth of the continental
rise, mainly by turbidites
between reefs (e.g. Seymen, 1975), continued in the eastern Pontides. The
area of the Rhodope-Pontide
Fragment between western Thrace and Zonguldak appears to have been emergent during this time period, probably
because of the Palaeo-Tethyan
deformations.
Upper Cretaceous sits with
angular unconformity
and a basal conglomerate
on the deformed Triassic
lithologies
of the Kocaeli (Bithnian)
peninsula (Abdiisselamoglu,
197’7). In
the eastern part of the eastern Pontides an as yet enigmatic “sill event” produced extensive tholeiitic basaltic sills (Lower Basic Series: see Akin, 1979)
- / ANATOLIDE
Fig. 6D. Palaeotectonic
map of the Iate
-TALlRIDE
-early
Jurassic
Cretaceous.
For legend see p, 206.
that invaded Appian-A~t~an
rocks. Whether this isolated event marks an
initial attempt at starting a subduction zone south of the eastern Pontides is
unclearir,
The Sakarya Continent was the site of neritic carbonate deposition (Bilecik Limestone: see Altinli, 1973a; Saner, 1977) during the late Jurassic. The
conditions changed rather rapidly to a pelagic environment during the early
Cretaceous and deeper water cherty carbonates (So@kCam Limestone,
Altinli, 1973a) were laid down. The submarine relief must have increased
during this time as evidenced by extensive slump deposits in this pelagic
milieu. The reason why the Sakarya Continent suddenly foundered is unclear,
but if it can be considered that the westerly extension of the Paikon “Ridge”,
the isostatic loading of the latter by the ~o-~e~en~~ ophiohte nappes during
216
the late Jurassic-early
Cretaceous abduction
event (Mercier et al., 1975;
Jacobshagen et al., 1976) may have extended its effects westward onto the
Sakarya Continent. Whether this is the case or not, we note here the remarkable temporal correlation between the two event,s.
On the Anatolide-Tauride
Platform also, at least two pelagic domains
made their appearance during the late Jurassic-early
Cretaceous. Although
neritic carbonate sedimentation
continued,
on most of the platform, pelagicsediment-filled
elongated troughs disrupted its continuity.
One such
basin may have separated the Menderes Massif from the Bey Daglari autochthonous domain and may represent an easterly continuation
of the Ionian
trough of the Hellenides (Gutnic et al., 1979). The number, geometry and
evolution of these pelagic “troughs” on the Anatolide-Tauride
Platform are
as yet not clear, however, (see fig. 26 in Gutnic et al., 1979), and that is why
we have not included them in our Fig. 6D.
Late Jurassic-early
Cretaceous
is also the time in which most of the
preserved Neo-Tethyan ophiolites of Turkey were produced. This is the same
picture as in the Alps and the eastern European Tethyan ranges (Dewey et
al., 1973) and similar to the situation there, is accompanied by the pelagization of parts of the previously neritic bordering carbonate platforms. The
coincidence
of the appearance
of ophiolites in the preserved record (not
necessarily their actual intial appearance) and the rather sudden foundering
of parts of the bordering carbonate platforms constitutes one of the current
enigmas of tectonics. It is difficult to ascribe this foundering to lithospheric
stretching alone because it occurs long after the continental rifting process
has ceased and synchronously
with the first direct evidence of spreading (not
necessarily first spreading) in the adjacent ocean. Most of the stretching of
the continental
crust generally occurs during the initial stages of rifting and
there is little extension in Atlantic-type
continental margins after plate accretion begins (e.g. Montadert et al., 1979). Initial rifting of the Neo-Tethyan
oceans in Turkey occurred during the late Triassic-earliest
Jurassic, whereas
the appearance of the first preserved ophiolites and widespread foundering
of previously neritic carbonate platforms were delayed until the late Jurassic
and/or early Cretaceous. A thermal explanation
is also difficult to devise,
because the agedependent
cooling and subsidence curve of Atlantic-type
continental
margins is close to exponential
and most of the subsidence
occurs shortly after rifting takes place (Hays and Pitman, 1973; Pitman,
1978). No model as yet satisfactorily explains this delayed rapid subsidence,
a situation very similar to the problems encountered
in modelling rift-related
basins such as the North Sea (see Burke, in
intracontinental
“epeirogenic”
press).
LATE CRETACEOUS-PALAEOCENE
EVENTS
(Fig. 6E)
During the earliest late Cretaceous clear evidence of subduction activity
appears all along the Pontides. According to the Africa-Europe
relative mo-
217
-
SOUTHERN
,,,,,
,,,
BRANC
F
L
Fig. 6E. Palaeotectonic
map of the late Cretaceous-Palaeocene.
For legend see p. 206.
tion paths of Pitman and Talwani (1972) and Dewey et al. (1973) there
should have been a considerable amount of oblique north-south convergence
between Europe and Africa between 148 m.y. ago (middle Kimmeridgian)
and 80 m.y. ago (Santonian). Until the Cenomanian-Turonian
(90-100
m.y. ago), however, we see almost no evidence of convergence anywhere in
Turkey with the possible exception of the localized pre-late Cretaceous folding event south of the Munzur Mountains and the enigmatic early Cretaceous
sills (?initial subduction) of the eastern Pontides. According to the AfricaEurope relative motion path of Biju-Duval et al. (1977), on the other hand,
the main Africa-Europe
convergence began during the latest Cretaceous
(76-68 m.y. ago), about 20 m.y. later than the local geological data indicate.
Following a brief episode of overall subsidence and pelagization during the
early/late Cretaceous transition (Fig. 5, columns 5 and 6), a strong magmatism
involving tholeiitic basalts, rhyolites and extensive tuff deposits began on the
eastern Pontides. These volcanics locally sit on the underlying pelagic limestones with a slight unconformity. Later in the Cretaceous, during the Campanian--Maastrichtian,
extensive volcanogenic flysch deposits began to
accumulate, particularly in the southern parts of the Pontides. The initial
intrusions of the Rize granitic pluton began 94 m.y. ago (Gedikoglu, 1978).
218
The late Cretaceous volcanic axis on the Rhodope-Pontide
Fragment in
Turkey swung northwestwards
to the west of Samsun and closely followed
the present Black Sea coast of the western Pontides. In a few places. such as
to the north of Istanbuf there are onshore outcrops of la&> Cretaceous an’related volcanic rocks. By and large, the presently exposed western portion of
the Rhodope-Fontide
Fragment was in a fore-arc position during the late Crctaeeous, in which thick flysch sequences of Turonian~enomanian
(Cenomanian wildflysch) age (Ketin, 1955) were deposited. Later during the Maastrichtian and the ?Palaeocene the flysch basin shallowed considerably
and
sedimentation
switched to biogenic limestone deposition (Ketin, 1955).
As the magmatic arc was developing on top of the Rhodope--Pontide
Fragment a melange wedge began growing in front of it. It is predominantly
of late Cretaceous age, but in places appears to extend into the Palaeocene
(Tokay, 1973). Blueschists are seen to be associated with this melange
locally along the strike of the Pontides.
During the late Cretaceous the Black Sea began opening behind the Rhodope-Pontide
island arc (Letouzey
et al., 1977). This opening event was
probably connected with that of the Srednogorie Province of the Balkanides,
where pelagic sediments composed of limestones, radiolarites and “<*ouches
and submarine volcanics of late Cretaceous age were
rouges”, turbidites,
deposited. Hsi.i et al. (1977) pointed out the great similarity between the setting and the contents
of the Srednogorie
province and modern marginal
basins and argued that it opened as a marginal basin behind the Rhodope--Pontide Fragment. During the Palaeocene-Eocene
the Srednogorie portion
of the marginal basin closed and the Emine Flysch was deposited in a forrdeep in front of north-vergent
thrusts (Hsu et al., 1977), while the Black Sea
portion remained open and subsided further, perhaps as a result of the continued cooling of its oceanic substratum (Letouzey et al., 1977). Because of
its short lifespan we did not include the Srednogorie portion of the greater
Black Sea marginal basin in our reconstruction
in Fig. 6E.
On the Sakarya Continent,
intensive talc-alkalic volcanism began during
the Turonian with the volcanics of the Vezirhan Formation (Altinli, 1973)
(Fig. 5, column 4). In the northern part of the continent this volcanism is
evidenced by numerous tuff interc~ations
in the red pelagic limestones of
the same formation (Altinli, 1973a). To the south of the Sakarya Continent,
an ophiolitic melange wedge developed between the Turonian and the Campanian and was overlain during the late Maastrichtian-Palaeocene
by shallow
water sediments in a regressive facies (Saner, 1978). The Intra-Pontide ocean
closed sometime between the Palaeocene and the Lutetian, because piantbearing Lutetian sediments unconformably
cover its suture northwest of
Ankara (Tokay, 1973).
The Pontides (Rhodope-Pontide
Fragment and Sakarya ~~ontinent) are
one of the few places in the world, where unequivocal
evidence exists to
show that subduction
began at the ocean-continent
contact zone and the
associated arc has been constructed
directly above the sedimentary prism of
219
the previous Atlantic-type
continental
margin (Yilmaz, in prep.; see Dewey,
1969a). On the Sakarya Continent the magmatic front marched southwards
into the area of the melange prism during the Palaeocene and migrated back
to the north for a brief time during the early Eocene following the closure of
the Intra-Pontide
ocean.
Throughout
the Anatolide-Tauride
Platform and the northern portion of
the Arabian Platform, the late Cretaceous is a time of extensive, preterminal
ophiolite abduction
(Ricou, 1971; Ricou et al., 1975). Previously, all these
events were thought to have been coeval, but new data indicate that obduction onto the Arabian Platform may have begun somewhat earlier than onto
the Anatolide-Tauride
Platform.
Particularly
the western portion of the
Anatolide-Tauride
Platform began subsiding en masse during the Maastrichtian as indicated by the invasion of nearly all previously neritic domains by
pelagic sediments (Delaune-Mayere
et al., 1977; Gutnic et al., 1979, particularly fig. 26). During the Campanian-Maastrichtian
the Bozkir Ophiolite
Nappe (named after ozgiil’s (1976) Bozkir Unit; the “Peridotite Nappe” of
the Lycian Nappes: Bernoulli et al., 1974) began climbing onto the Anatolide-Tauride
Platform. This is evidenced by the deposition of Maastrichtian
flysch and ophiolitic olistostromes
on top of the northern continental shelf
of the Anatolide-Tauride
Platform, which was actively subsiding under the
weight of the advancing ophiolite
sheet(s) (e.g. the Kijycegiz region, De
Graciansky,
1972; the Haymana Basin, N. Goriir, personal communication,
1980). Sparse blueschist facies metamorphism
beneath the advancing thrust
sheets, such as those of the Emirdag-Kiitakya-Eskigehir-Balikesir
belt,
where such metamorphics
are found in an allochthonous
position above the
Anatolide-Tauride
Platform (Bingiil, 1978) are probably related to the ophiolite emplacement.
The abduction of the Bozkir Nappe o&m-red all along
the northern margin of the Anatolide-Tauride
Platform during the Campanian-Maastrichtian
(Bergougnan,
1975; Di.irr, 1975; Ricou et al., 1975;
Gzgul, 1976; ijzgul et al., 1978) and marks the beginning of the Neo-Tethyan
deformations
here. The Maastrichtian events have not carried the ophiolitic
allochthons very far onto the platform, as shown by the continuous platform
sedimentation
along its central axis from the Menderes Massif-Bey
Daglari
area in the west to the Sariz region in the east (Fig. 5, columns 7-10, 13),
contrary
to the interpretations
of the Orsay Group (Ricou et al., 1975;
Delaune-Mayere
et al., 1977; Gutnic et al., 1979), who would like to extend
the Bozkir Nappe to Cyprus and southeastern Turkey in the south.
Ophiolite
abduction
onto the Bitlis-Poturge
continental
fragment also
occurred at about the same time. On top of the Upper Cretaceous section of
the outer envelope, the ophiolite nappe rests above a zone of intense cataclasis (approximately
500 m thick) and a thin ?slice of (O-30 m) metamorphic aureole
containing
albiteactinolite-epidote-chlorite
assemblages
(Fig. 5, column 15) ?Palaeocene-Eocene
molasse unconformably
covers the
ophiolites in the western part of the Bitlis Massif (Yilmax, 1978; in prep.),
Yilmaz (in prep.) ascribes much of the metamorphism
of the Malatya-
220
Keban, Poturge and Bitlis massifs to this abduction
event. l‘he metamorphism of these massifs is of burial type and may have occurred as a result of
the descent of the Bitlis-Piitiirge
continent int.o progressively hotter regions
as it was tucked beneath the overriding oceanic lithosphere, similar to the
situation in the western Fleur de Lys in Newfoundland
(Williams, 1977). ‘The
ophiolite emplacement in the Aladag region (Tekeli, 1978), along strike from
the Bitlis-Pijtiirge
area to the west-southwest,
is probably a part of t,he same
event.
Immediately
after or synchronously
with this abduction event, a southdipping subduction
zone beneath the ophiolite-laden
Bitlis-Poturge
continent began consuming the floor of the Inner Tauride Ocean. The evidence
for this subduction
zone comes largely from the Yiiksekova arc lithologies
that include pyroclastics, pillow basal&, diorites of various compositions and
possibly some andesites (Hemp~n
and Savci, in press). Near Gijzeli, to the
west of Lake Hazar, the plutonic equivalents of the Yiiksekova arc magma
tism include large granodiorite
plutons that are unconformably
overlain by
Palaeocene lithologies. In the Bolkardag region to the south of the western
extremity
of the Inner Tauride ocean, arc magmatism and associated ore
mineralization
lasted until the late Palaeocene-?early
Eocene (Calapkulu,
1978).
The Alanya Massif overrode
the small part of the Pamphylian
basin
between itself and the Anatolide-‘I’auride
Platform and terminated its sedimentation during the late Cretaceous. N. iizgiil (personal communication.
1978) indicates that there might be evidence for an incipient arc magmatism of late Cretaceous-Palaeocene
age in the Alanya ILlassif that could be
responsible for its metamorphism,
but this remains to be demonstrated.
The southernmost
of the late Cretaceous abduction events in Turkey and
the eastern part of the eastern Mediterranean
was the Senonian ophiolite
emplacement
that extended all the way from Cyprus to the Turkish-Iraqi
border.
In Cyprus the Mammonia Nappes (Lapierre, 19’75) probably originated as
a subduction-related
melange complex in front of the Kyrenia magmatic arc
(Baroz, 1979) and was later thrust over, together with the Troodos complex
(Gass, 1968; Biju-Duval et al., 1976) on top of African oceanic or stretched
and thinned continental
lithosphere.
Our Fig. 6E shows the penultimate
imbrication
stage to emphasize the original north--south
thrusting of the
Mammonia Nappes as origin~ly proposed by Lapierre (19751 as opposed to
the south-north
thrusting suggested by Robertson and Woodcock (19791,
evidence for which seems flimsier than for southward thrusting. Along the
northern periphery of the Arabian shield, abduction may have begun as early
as earlier Campanian (Al Maleh, 1976) in the western part, but the bulk of
the ophiolites seem to have been emplaced during the late Campanian-early
Ma~trichtian
(Sunguriu, 1974) which is also the age bracket of the Kastel
fiysch that developed in front of the advancing alloehthons.
The abduction
onto the Arabian Platform may have been a result of the collision between
221
the latter and the Bitlis-Poturge continent as in several places the massifs
and the supposed correlatives of the ophiolite nappes (the Guleman ophiolite,
Perinqek, 1979) have a common latest Ma~t~ch~~
sedimentary cover.
However, the correlation between the Senonian nappes of southeastern Anatolia and the Guleman ophiolite is not yet well established. We adopt the collision hypothesis, because it seems to be the best one currently available.
The opening of the Maden marginal sea complexes probably began during
the latest Ma~trichtian, i.e. shortly after the south-dipping subduction had
been initiated north of the now combined Bitlis-Pijtiirge and Arabian continent.
The Simaki “Flysch” (Perincek, 1979), a late Maastrichtian-Palaeocene
sequence that begins with coarse elastics on top of the Bitlis-Poturge
metamorphics and Guleman ophiolites and rapidly changes to progressively
deepening shallow water limestone litholo~es and some turbidites probably
mark the initial rifting and beginning subsidence of the Maden marginal sea
complexes (Fig. 5, column 16); there is a small amount of volcanics in the
Middle Palaeocene section that may have been related to the rifting event.
The Campanian-Maastrichtian
ophiolites in the east of Lake Van (0. Sungurlu, personal communication, 1979) may represent continued spreading
activity in the Neo-~ethyan ocean floor here and imply a complex plate
boundary geometry.
The late Cretaceous was a time of revolution in the Neo-Tethyan tectonic
development of the Alpides in general and this has been so in Turkey as well.
It marked the beginning of the convergent regime here “at all fronts” and
was particularly characterized by the emplacement of spectacular ophiolite
nappes of large dimensions, These nappes moved onto extensive carbonate
platforms that began subsiding en masse synchronously with the onset
of abduction to become sites of either pelagic or flysch-olistostrome
deposition depending on their distance to the advancing ophiolite sheets. Along the
southern borders of the Rhodope-Pontide
Fragment and the Sakarya Continent ophiolitic melange was overthrust northwards onto the arc terrains
(Bergougnan, 1975; Fourquin, 1975; mostly our observations). The magnitude of this overthrusting is small, however, and nowhere does it involve
large, coherent slabs of oceanic lithosphere as is the case on the passive margins to the south. We interpret this thrusting as a retrocharriage phenomenon
at the contact between the arc massif and the accretionary melange wedge.
As the melange wedge grows oceanward, the wedge-shaped thrust packages
rotate to steepen and eventually to overturn (Karig, 1974). This overturning,
combined with the increased gravitational potential of the thickened melange
wedge begins driving thrust sheets at the back of the wedge backwards onto
the arc massif (Hamilton, 1979). Weber (1978, fig, 8) has recently described
a very beautiful fossil example of this phenomenon from the Hohensteiner
Sattel from the Rhenoher~ynian zone near Giessen, W. Germany. This
“simultaneous” ophiolite thrusting onto the Pontides and the AnatolideTauride Platform during the late Cretaceous and the existence of shallow-
water and even terrestrial deposits on thickened melange wedges during the
Palaeocene has been incorrectly interpreted in the past to indicate a terminal
closure of the northern branch of Neo-Tethys in Turkey. As we shall see in
the next section, however, that closure actually occurred during the latest
Palaeocene-early
Eocene times,
EARLY
TO MIDDLE
EOCENE
EVENTS
(Fig. 6F)
During the ?latest Palaeocene-early
Eocene the Anatolide-Tauride
Platform collided with the Pontides. Immediately
following this collision largescale internal deformation
of the Anatolide-Tau~de
Platform began that
was synchronous
with the extensive r~troch~iage
development
in the Pontides.
Arc magmatism continued throughout
the Pontides during the early and
medial Eocene as evidenced by extensive talc-alkalic andesitic lavas, pyroelastics and volcanogenic flysch (Fig. 5, columns 4-6), while the deformed
Pontide-Anatolide
suture zone was overlain by shallow water and terrestrial, largely medial Eocene deposits along the western sector of the suture
(Tokay, 1973).
The first evidence of the beginning of the internal imbrication
of the
Anatolide-Tauride
Platform is provided by Parejas’ (1943b) “phase anatolienne” which corresponds to the burial of the central Anatolian crystalline
massifs (~enderes and Kirsehir), the Anatolides of Ketin (1966), under the
advancing Bozkir Nappes (Lycian, Beysehir Hoyran and Hadim Nappes plus
Fig. 6F. Palaeotectonic map of the early-middle
Eocene. For legend see p. 206.
223
the thrust sheets that once lay atop the Kirsehir Massif; &q$il, 1976) (Gutnit et al., 1979). Ypresian and Lutetian sediments everywhere lie unconformably on the structures of this phase, which also corresponds to the beginning
metamorphism
that eventually formed the Anatolides.
There has been a great diversity of opinion as to the age of the Anatolides.
Earlier during this century workers such as Kober (1921) interpreted them as
pre-Alpine
median massifs (“Zwischengebirge”)
that had influenced
the
Alpine development
of the Turkish orogen, a view adopted, among others,
by Leuchs (1943) and more recently by Brinkmann (1976). Another extreme
is represented
by Ketin’s view (1959, 1966), who considered the late Cretaceous-early
Tertiary deformations
as the first erogenic event that had ever
affected this area. Some other workers have taken more compromising views
(e.g. Izdar, 1975). Recent work on the Anatolides, particularly on the Menderes Massif (Boray et al., 1973; Diirr, 1975; Diirr et al., 1978; Gutnic et al.,
1979), has shown that the massifs have a late Precambrian (Pan AfricanBaykalian) continental
basement on which Palaeozoic, Mesozoic and early
Tertiary (Lower Eocene inclusive) sediments were deposited without any
major break. The same story holds true for the Bitlis and Potiirge massifs
until the late Cretaceous
ophiofite event. The Anatolides metamorphism
as a result of their burial beneath the Boxkir Nappes during the Anatolian
Phase (Diirr, 1975; SengBr, 1979b). R. Akkijk (personal communication,
1979) has shown that the Menderes Massif was probably buried beneath
some 6 km of overburden.
The events in the Kiqehir Massif may have happened somewhat earlier
than in the case of the Menderes Massif. Both in its northern (Ketin, 1956)
and southern extremities (the Nidge Massif, GGnciioglu, 1977) the metamorphics of the massif are covered by unmetamorphosed
early and medial
Eocene flysch deposits. Late Eocene andesitic volcanics also sit on the metamorphics of the Kirqehir Massif at various places.
While the main body of the Anatolide-Tauride
Platform was being internally imbricated
and the initial burial metamorphism
of the Anatolides
began, the Maden and Ci.ing-iis basins in southeastern
Anatolia were about to
reach their maximum development.
In both of them deep-sea sediments
(pelagic limestones and radiolarian cherts) were being deposited accompanied
by a strong mafic volcanism mainly in the form of pillow lavas and some
turbiditic activity (Perirqek, 1979 and our observations).
In the meantime,
the Yiiksekova arc had ceased its activity during the early Palaeocene and the
anti-clockwise
rotation
of the Bitlis-Poturge
Massif fragments was being
accommodated
by a north-dipping
subduction zone beneath the eastern half
of the Anatolide-Tauride
Platform, The earliest evidence for this northward
subduction comes from the Palaeocene talc-alkalic volcanics of the Ulukisla
Basin (Calapkulu, 1978). The blueschist metamorphics
of the Bolkardag area
(Fig. 6E) are associated with the late Cretaceous ophiolites and have formed
sometime between the Campanian and the Lutetian (Calapkulu, 1978). The
Palaeocene-early
Eocene times also witnessed a strong flysch and wildflysch
224
deposition in the Ulukigla area, the western-most portion of the lnner I’auride
Ocean, synchronously
with the development
of the East Anatolian Accretionary Complex to its maximum size. During the same time the neritic Midyat Limestones
were being deposited on the southern shelf of the CiingQ
Basin, indicating the return of the quiet shelf conditions
to the northern
margin of the Arabian Platform (Sungurlu, 1974; Perincek, 1979).
The Alanya Massif with its cushion of Antalya Nappes was emplaced into
the future Isparta Angle (Courbure
d’Isparta, Blumenthal,
1963; Monod,
1976) during the late Palaeocene-early
Eocene (Brunn et al., 1971; DelauneMayere et al., 1977).
LATE
EOCENE
TO EARLY
MIOCENE
EVENTS
(Fig. 66)
During the late Eocene to early Miocene interval the general north-south
tightening of the Turkish orogen continued
while the Anatolides were uplifted and unroofed. Late Eocene in the east (Kirsehir) and Oligocene in the
west (Menderes) cover most of the area of the crystalline massifs. However,
as the tightening continued and the Bozkir Nappes gradually marched into
their final resting places more and more continental
material was being fed
beneath the massifs, which helped their uplift and gave rise to deep crustal
partial melting
melting under the over-thickened
sectors. This anatectic
caused the later widespread silicic volcanism throughout the western Anatolia
and granitic plutonism
in the Aegean islands (Izdar, 1975; BingoI, 1976;
Diirr et al., 1978). During the late Eocene-Oligocene
the Beysehir-Hoyran
and the Hadim composite nappe systems reached their final destinations. This
was accompanied
by the southerly-imbrication
of the autochthonous
series
Fig. 6G. Palaeotectonic
map of the late Eocene
and early Miocene.
For legend see p. 206.
225
of the eastern flank of the Isparta Angle (Monod, 19’79). Before the deposition of the Oligocene molasse in the Kale-Tavas Basin (Fig. 2; Becker-Platen,
1970), the Lycian Nappes also advanced onto the autoch~onous Bey Daglari
region during the Kale Phase of Poisson (in Gutnic et al., 1979). The Oligocene Kale-‘I’avas molasse seals the thrust contacts between the Lycian nappes and the unroofed Menderes Massif that has a tectonic window position in
western Turkey (Diirr, 1975; Gutnic et al., 1979; Sengiir, 1979b).
In east-central and southeastern Anatolia the late Eocene saw the closure
of the Inner Tauride ocean and the Maden basin (Calapkulu, 1978; Perinqek,
1979). This event internally imbricated and interleaved the Yiiksekova,
Maden and the Bitlis-Poturge lithologies and the Eocene flysch fill of the
Inner Tauride ocean. The separation of the Malatya-Keban metamorphics
into a separate slice from the Bitlis-Pijtiirge metamorphics was probably a
rna~~ifes~t~on of this collision event which also deformed the western half
of the East Anatolian Accretionary Complex and essentially terminated its
growth, perhaps with the exception of the area to the southeast of the present Lake Van. Immediately following this collision, the C$ngii$ basin began
receiving large amounts of flysch sediments and olistostromes with blocks of
mainly Maden and Bitlis-Pijtiirge lithologies. Sandstone blocks in the $%ngiis olistostromes may well represent eroded pieces of reworked Inner Tauride flysch.
Either synchronously with or just before the terminal closure of the InnerTauride and the Maden basins a flare-up of andesitic volcanism occurred
along a wide belt from Yozgat to Kars (Yozgat-Sivas volcanics of Fig. 2 are
the western representatives of this event). We associate this volcanism with
the subduction of the Inner Tauride ocean beneath the region of volcanism.
The Palaeocene volcanism of the Uluki.$a area (Calapkulu, 1978) is here
considered as the initial appearance of the same event, which prematurely
terminated as a result of the terminal closure of the Inner Tauride ocean.
Following this closure the Africa-Eurasia convergence began to be taken
up along a sinuous, uniformly no~hdipping subduction zone beneath southern Turkey. This subduction zone was consuming very young ?oceanic lithosphere of the Qingiis basin along its eastern extent whereas south of central
and western Anatolia much older (?middle Mesozoic) lithosphere was being
subducted. This geometry might well be responsible for the Eocene initiation
of the left-lateral Ecem$ Fault (Fig, 2) that accentuated the sinuosity of the
subduction zone by acting as a trench-trench
transform separating the
extensional (western) segment of the arc from the compressional (eastern)
one -we have not indicated this geometry in Fig, 6G, as the details to confirm or reject this hypothesis are not available to us.
Within the Anatolides, continued uplift under ongoing north-south
shortening resulted in some retrocharriage development that is evidenced in the
Menderes Massif by ductile, south-dipp~g shear zones of Miocene age (E.
Izdar, personal communication, 1977), whereas in the Kirgehir Massif, near
Kaman, cold basement slices overthrust Eocene and some Neogene sediments
226
with a northerly
vergence.
(I. Seymen
and F. Oktay, personal
communication, 1977).
The Rize Pluton in the eastern Pontides
completed
its intrusion
during the
late Eocene (Cogulu,
1975).
EARLY
TO MIDDLE
MIOCENE
EVENTS
(Fig. 6H)
During the early to middle Miocene interval
the Ciingiis basin terminally
closed and the Arabia-Eurasia
collision
started along the Bitlis Suture while
the Lice molasse
(Sungurlu,
1974; Perincek,
1979) was being deposited
in
the foredeep
that formed
on the Arabian
Platform
as a result of this collision. Very shortly
thereafter
terminal
collision
also occurred
along the Zagros suture,
and Arabia
became
weled to Eurasia
along the Bitlis-Zagros
suture zone (Dewey
et al., 1973; Sengor,
1979b;
$engor
and Kidd, 1979,
$engiir et al., 1979). This collision
had very profound
effects on the overall
tectonics
of Turkey
and has essentially
governed
its post-Seravallian
evolution.
As the north-south
shortening
across
eastern
Turkey
continued
between
the converging
jaws of Eurasia
and Arabia the relatively
soft and
irresistant
East Anatolian
Accretionary
Complex
took up much
of the
initial
post-collision
convergence
by shortening
and thickening.
However,
the rapidly
rising elevations
made it eventually
more economic
to wedge out
of the way a considerable
piece of Turkey,
roughly
coincident,
particularly
in the east, with the original
outlines
of the Anatolide/Tauride
platform,
onto the easily subductable
eastern
Mediterranean
floor (McKenzie,
1972;
qengor,
1979b; 1980; qengiir and Kidd, 1979). Thus, the North and the East
Anatolian
transform
faults, and with them the Anatolian
Plate, originated
(Sengor, 1979b).
I
ANATOLldN
N-S extension
here
begins 10: p.st-Torton v4_ r:?t-Tortc
Lice
-- ‘;‘s
Molasse \---~>,
ARABIAN
Fig. 6H. Palaeotectonic
map of the middle
Miocene-Pliocene.
PLATE
--.
For legend see p. 206.
227
Early Miocene was also the time of the final emplacement
of the Lycian
Nappes, which overrode a Burdigalian
molasse. The associated thrust contacts are sealed by Tortonian
conglomerates
(Altinli, 1945; Delaune-Mayere
et al., 1977). As the fill of the Kale-Tavas
basin seals the Oligocene thrust
contacts
at the back of the Lycian Nappes, the Miocene thrusts must pass
below this seal and probably represent a new crustal imbrication
of the western part of the Anatolide-Tauride
Platform with a minimum thrust transport of 100 km (Gutnic et al., 1979).
The Adanasilicia
basin, an intracontinental
basin that formed as a result
of the incompatibility
problems arising in a continental
FFF triple junction
near Marag (Sengor et al., in press) also formed during this time. The East
Anatolian
and the Miocene Dead Sea transform
faults meet at this triple
junction,
and because not all of the left-lateral offset can be accommodated
by shortening across the Bitlis suture zone, the dagger-shaped Adana-Cilicia
basin has to extend its area in a roughly east-west
direction.
During the late middle Miocene (Tortonian)
the eastern flank of the Isparta
angle overrode
the western flank at the northern
tip of the structure
(Delaune-Mayere
et al., 1977). Sengor and Dewey (in press) interpreted this
tightening of the Isparta Angle as a result of the westward movement of the
Anatolian
plate. The early Pliocene sediments unconformably
overlie this
thrust (the Aksu thrust). Since that time the consuming and transform plate
boundaries in the eastern Mediterranean
appear to have assumed their present geometry.
The Aegean extensional
regime also began at the same time due to a relief
of the east-west
shortening caused by the abrupt south-westerly
bend in the
North Anatolian transform fault west of the Sea of Marinara by north-south
extension (Dewey and Sengor, 1979; Sengiir and Dewey, in press).
PLIOCENE TO PRESENT
(Fig. 61)
Continued
convergence
and thickening
in the Turkish-Iranian
high plateau resulted in the widespread Plio-Quaternary
Tibetan-type
volcanism on
the plateau (Sengijr and Kidd, 1979), while the Border Folds formed on the
Arabian Platform as a foreland fold-thrust belt associated with the Miocene
collision
along the Bitlis-Zagros
suture (Ketin, 1966).
This collision also
gave rise to two impactogens
on the Arabian Platform, namely the Akcakale
Graben (0. Sungurlu, personal communication,
1979) and the Karacalidag
shield volcano (Qengiir et al., 1978).
While the Aegean extensional
regime
continued its activity, the strike-slip dominated ova regime became established
in central Anatolia (Sengiir, 1979b,
1980). The active tectonics of Turkey
represents
the continuation
of the Plio-Quaternary
regime and has been
described
in detail in the recent papers by Sengiir (1978,
197913, 1980),
Dewey and Sengiir (1979) and Qengijr and Kidd (1979).
A detailed discussion of the post-Tortonian
tectonics of Turkey is therefore omitted in this
paper.
Fig. 61. Palaeotectonic
DISCUSSION
map of the Pliocene-Present.
For legend see p. 206.
AND CONCLUSIONS
We have traced the outlines of the plate tectonic evolution of Turkey and
its immediately
neighbouring
areas since the Permian. This evolution has
been, in the simplest terms, the resul+ ,,f t!z destruction
of Palaeo-Tethys
and the birth and the subsequent demise of Neo-Tethys. In detail, however,
the sequence of events has been much more complicated, characterized
by a
number of mini-oceans that have repeatedly opened and closed throughout
the time interval considered. What is particularly instructive is that although
the very broadest outlines of Tethyan geology can be deduced from the
evolution
of the world plate mosaic, its details are nearly impossible to
deduce from the motions of the bounding mega-continents
of Laurasia and
Gondwana-Land.
For example, when head-on convergence
between Africa
and Europe was going on during the middle Eocene (44 m.y. ago; Biju-Duval
et al., 1977) the Maden and the CiingQ basins were experiencing the zenith
of their extensional history. Similarly, as the two continents were receding
with respect to each other during the early and middle Jurassic, the Pontides
were experiencing intense subduction and later collision tectonics. Throughout the Tethyan history plate boundaries were constantly shifting their position and often changing character, particularly during the late Cretaceous-late Eocene interval (Figs, 6E-G), when subduction zones were constantly
“jumping around” in Turkey to continually accommodate
the ongoing convergence. Some of these jumps, for example that which occurred during the
late Eocene and moved the subduction activity from the northern margin of
the Inner-Tau~de
ocean to the northern margin of the ~~n~~ Basin resulted
from continental
collisions, whereas others, such as the switch from the late
Cretaceous-Palaeocene
south-dipping subduction along the southern margin
229
of the Inner Tauride ocean, to Palaeocene-late
Eocene north-dipping subduction along its northern margin with coeval “quitting”
of the former, occurred for less obvious reasons. Modern examples of such “quitting” subduction
zones are also present, the best known example being perhaps the subduction zone of the Palawan Trench in the Phillipines that quit during the Miocene (Hamilton,
1979),
and are probably the consequence
of local plate
boundary reorganization.
As a result of this rapid change in plate boundary
position and geometry,
the life histories of the platelets bounded by them
become extremely
complicated.
In many places the geologic record indicates
decidedly episodic deformation
although the overall displacement
between
plates was more or less continuous as a result of such abandonment
and reestablishment
of plate boundaries, particularly
of subduction
zones. Dewey
(1975) has shown the amazing level of complexity
in tectonic evolution that
can be generated as a result of the continuous evolution of only three plates.
In regions where more plates interact, the continuous
shifting of the instantaneous rotation poles could create bewildering complexities,
the record of
which may be completely
obliterated
by subduction or by deformation
during a subsequent
continental
collision.
During the Permian to present tectonic evolution of Turkey we have observed a series of events that rifted,
combined and re-rifted, along different lines, the microcontinents
within the
Tethyan domain. Some of these “microcontinents”
were of enormous size
and extended well-beyond
the boundaries of Turkey, such as the Cimmerian
Continent,
which extends from northern Greece and Bulgaria to the east of
Tibet (Sengor, 1979a).
Others were much smaller such as the Alanya Massif
and the Bitlis-Piitiirge
Continent.
Another
interesting
observation
is the “triggering”
effects
that plate
boundaries seem to have on each other. We have seen that the Neo-Tethys in
Turkey began opening, at least partly, as a series of back-arc basins over the
Palaeo-Tethyan
subduction
zone. Similarly
Maden and Ciingiig basins are
marginal seas related to the Yiiksekova subduction
zone, whereas the Black
Sea is a marginal basin triggered by the Cretaceous subduction zone that dipped beneath the Rhodope-Pontide
Fragment.
Such reciprocal
effects of
plate boundaries on each other are very widespread particularly in the history
of the Tethyan belt (Qengor, in press) and, as the prediction of the tectonic
style of any arc system depends on a knowledge of the age of the underthrusting oceanic lithosphere
(Molnar and Atwater, 1978) and of the relative motion of the overriding plate with respect to an inert asthenosphere
reference
frame (Wu, 1978; Dewey, in press), it is almost impossible to foresee where
and when they should occur due to the inherent uncertainties
of field data
and the self-destructive
character of plate tectonics.
In Turkey, there is also an important change in tectonic style of collision
along strike., The profiles in Figs. 3A and B were drawn to show this difference and Figs. 3A’ and 3B’ to emphasize it schematically.
In western Turkey the structure is dominated, south of the Izmir-Ankara
suture, by far-travelled composite nappe systems that become progressively
younger in the direction of transport (suture progradation of Roeder, 1979)
and tectonic windows into rising metamorphic
complexes (the Nlenderes
Massif) that represent the parautochthon
(inner part of the Anatolide--Tauride Platform) on which the nappes (Bozkir Unit) rest. The rise of the Menderes Massif began about 15 m.y. after its burial and continued until the
extensional regime of western Anatolia interfered in post-Tortonian
times.
Synchronously
with the uplift of the Menderes Massif, south-directed
thrustprogradation
in the Lycian Taurus and north-directed
rktrocharriage
in the
western Pontides continued.
This geometry and evolution seems to be the
result of the continuous imbrication of the Anatolide-Tau~de
Platform following the terminal closure of the Izmir-Ankara
ocean. During the Anatolian
phase, a part of the platform gets buried beneath allochthons and becomes
metamorphosed.
The hot, ductile core of the metamorphic
complex begins
rising in a convergent environment
and starts rotating the suture zone into
a vertical position. As the convergence continues the suture zone overturns
and the hinterland that was originally the overriding plate begins inserting
itself “into” the side of the ductile metamorphic
uplift as a wedge. This is
expressed at the surface as late rktrocharriage (Fig. 3A’). This wedge initiates
a separation between the upper and the lower continental crust which propagates to the surface in the south as listric thrusts that involve only the upper
part of the crust and eventually root into the hot metamorphic
core complex, the active examples of which may be expressed as low-velocity zones in
the cores of orogens (e.g. beneath the Central Alps: Miller et al., 1976; Hsii,
1979). While the upper crust is being piled up in the form of successive
thrust sheets (the late Eocene, “pre-Kale-‘I’avas--Oligocene”
and the postBurdigalian thrusting episodes in the Lycian Taurus), the lower crust can be
removed by subduction (Molnar and Gray, 1979).
In eastern Turkey, along the profile B, the style is dominated by steeply
to moderately
dipping thrusts and numerous ophiolitic sutures. Along this
profile there is no equivalent of the Menderes-type
metamorphic
uplifts. A
glance at the geological evolution of this part of Turkey reveals the essence
of this style. Here, much of the convergence until the Miocene was accommodated by the subductive removal of the floors of the numerous minioceans that separated a number of microcontinents,
Intra-continen~l
imbrication, the essence of the flat nappe structures and the Menderes-type metamorphic uplifts, has been absent in eastern Anatolia until the rMiocene. Continuous suturing of microcontinents
has created there a series of moderately
to steeply dipping suture zones with abundant ophiolites. The limiting case
of the latter tectonic style is to squash a melange wedge between two continents (see Dewey, 1969b for an excellent example of such a situation),
which has been the case for the East Anatolian Accretionary
Complex still
farther east.
Another important aspect of the East Anatolian Accretionary
Complex is
that it represents a “continental
hole” created as a result of the collision of
four continental
objects that did not have perfectly matching margins and
231
ACCRETIONARY
MATERIAL
BULAWAYAN GREENSTONE
BELT RHODESIA
0
Km
20
LESSER
CAUCASUS
NORTH
CHINA
PLATFORM
PLATFORM
INDIAN
PLATFORM
Fig. 7. Sketch maps showing three “deer’s head” shaped, accretionary
material-filled
continental holes that resulted from the collision of more than two continental
objects. B
and C were drawn to the same scale. A is drawn after Burke et al. (1976)
and C is from
the geological map of China (1976).
Discussion in text.
filled with oceanic island arcs and melange complexes. The final collision of
it with Arabia further deformed it internally and “consolidated”
it by magmatic intrusions (see Sengor and Kidd, 1979). Figure 7 shows a comparison
of the East Anatolian Accretionary
Complex with two other possible analogues, the Precambrian greenstone belt of Bulawaya and the Sungpan-Kantze
system of China. Burke et al. (1976)
pointed out that the Bulawayan greenstone belt probably originated as a result of the collision of small island arcs
today represented
by the granodioritic
terrain that surrounds the triangular
greenstone belt. Continental crust has grown, they argued, by the coalescence
of such arc systems by repeated collisions during the Archaean. Similarly,
the Sungpan-Kantze
system is the “continental
hole” that resulted from the
imperfect closure of Palaeo-Tethys
among four continental
objects (Sengor,
232
in press). Its fill and overall geometry
have an amazing
resemblance
to the
East Anatolian
Accretionary
Complex,
the difference
being that the former
is about three times as large as the latter and about 150 m.y. older. Such
“deer’s
head”-shaped
accretionary
complexes,
caught
up in suture
knots
where continents
cannot
fill holes that result from irregular margin collision,
may be more widespread
than previously
recognized
and may contribute
substantially
to the areal growth of continents
as has obviously
been the case in
Turkey and China.
Although
extremely
complicated
and less well-exposed
than many other
erogenic
belts in the world, Turkey
offers perhaps
a unique
place for the
study of Tethyan
geology in particular
and collision
processes
in general. It
not only has a representative
selection
of nearly
all of the main Tethyan
palaeogeographic
elements,
but has an amazingly
rich assortment
of collision
related structures.
After we obtain a reasonably
clear idea of what the outlines of Tethyan
evolution
were, the next important
step will be to try to fit.
the Hercynian
deformations
of Turkey
into their proper
place within
the
overall geometry
of the Hercynian
orogen in Euro-Africa.
ACKNOWLEDGMENTS
We are grateful
to Mr. Esen Arpat, Dr. Ergiizer Bingol, Prof. Ihsan Ketin,
Mr. Necdet ijzgiil and Mr. Ozan Sungurlu
for discussions
and the unpublished
information
they have so generously
provided.
Discussions
with Professors
B. Clark Burchfiel,
Kevin Burke, J.F. Dewey and Drs. R. Akkok,
N. Goriir,
W.S.F. Kidd, J. Marcoux,
0. Monod, A. Poisson, L.-E. Ricou, 1. Seymen and
N. Yalqin were most helpful.
Many other colleagues
and friends at the Mineral Research
and Exploration
Institute
of Turkey
(the M.T.A.)
and the
Turkish
Petroleum
Co. were most helpful.
Albany graduate
students
working
in Turkey,
Messrs. Gary W. White,
Mark Hempton
and Miss July Dyer
kindly made much of their data available
to us. Miss D. Kelly of the departmental staff at SUNY patiently
typed this manuscript.
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